Middle Eocene to early Oligocene intrusions, widespread in the Ruby Mountains–East Humboldt Range metamorphic core complex, Nevada, USA, provide insights into a major Paleogene magmatic episode and its relation to tectonism in the northeastern Great Basin. These intrusions, well-exposed in upper Lamoille Canyon, range in composition from gabbro to leucomonzogranite. They form small plutons, sheets, and dikes that intrude the metamorphic and granitic infrastructure of the core complex. Two types of Paleogene monzogranite were recognized. The first is exemplified by two of the larger intrusive bodies, the Snow Lake Peak and Castle Lake intrusions, which occur as sheet-like bodies near and at the structural base of metamorphosed Neoproterozoic and Cambrian Prospect Mountain Quartzite where it is inverted above Cambrian and Ordovician marble of Verdi Peak in the Lamoille Canyon nappe. Swarms of dikes are associated with these intrusions. U-Pb (zircon) ages range ca. 40–33 Ma and typically display relatively simple and minor inheritance. The rocks have the lowest εHf (zircon) and εNd (whole rock) of any of the middle Cenozoic granites. The second type of monzogranite, Overlook type, typically occurs as thin, isolated dikes and leucosome-like bodies in Late Cretaceous granites of the infrastructure, with no obvious relationship to the large monzogranite bodies. Overlook-type monzogranite displays complex zircon inheritance, yields igneous ages ca. 37–32 Ma, and has εHf (zircon) and εNd (whole rock) identical to those of Late Cretaceous granites in the core complex. These isotopic and field data indicate that Overlook-type monzogranite formed in situ through anatexis of host Cretaceous granites. A pervasive thermal event was required to stimulate this crustal melting.

Gabbros from Lamoille Canyon and quartz diorite dated from 32 km away signal mantle-derived magmatism ca. 39–37 Ma (U-Pb, zircon) was a driver of crustal melting and hybridization. Eocene 40Ar/39Ar apparent ages on hornblende and biotite are consistent with syn- to post-magmatic extensional exhumation and decompression. Thus, the core complex provides a window into trans-crustal magmatism and insight into how such magmatism affected the Nevadaplano orogenic plateau. This Paleogene thermal pulse, which may relate to removal of the Farallon slab by delamination of mantle lithosphere, involved partial melting of the upper mantle and transfer of magma and heat to the Nevadaplano crust. Lower-crustal melting of Archean(?) to Paleoproterozoic rocks resulted in Snow Lake Peak–type magmas, and middle-crustal melting of granite in the infrastructure yielded Overlook-type magmas. This crustal magmatism, as exemplified by the intrusions in the core complex, likely played a role in destabilizing the Nevadaplano and its later collapse during middle Miocene extension. The Paleogene intrusions and associated structural features also provide insight into the evolution of the core complex through either the buoyant upwelling of a melt-rich diapir (gneiss-dome model) or buoyant upwelling of the melt-rich middle crust synchronous with a west-rooted mylonitic shear zone (extensional shear-zone model). We favor a hybrid that incorporates both models.

The relationship between the formation of metamorphic core complexes and magmatic activity has been a hotly debated topic. With the combination of detailed geologic mapping, geochronological studies, and geochemical and isotopic analyses, it is possible to address this subject in the context of the Ruby Mountains, northeastern Nevada, USA. This metamorphic core complex has been interpreted as recording Cenozoic, large-magnitude crustal extension, which exposed the middle crust in the hinterland of the Early Cretaceous–early Eocene Sevier orogenic belt (Coney, 1980; Armstrong, 1982; Snoke et al., 1990, 1997; DeCelles, 2004; Sullivan and Snoke, 2007; Fig. 1). However, recent studies of the Ruby Mountains–East Humboldt Range metamorphic core complex (Zuza et al., 2022; Zuza and Cao, 2023; Levy et al., 2023; Zuza and Dee, 2023) have challenged the large-magnitude extension model involving a rooted mylonitic simple-shear zone and associated upper-crustal detachment fault and, instead, have argued for buoyant doming of a melt-rich diapir with a mylonitic carapace characterized by general shear. A similar model of diapiric uplift of a granite-cored gneiss dome has also been suggested for the development of the Albion–Raft River–Grouse Creek metamorphic core complex in northwestern Utah, USA, and southern Idaho, USA (Konstantinou et al., 2012, 2013). A detailed understanding of the nature and timing of associated magmatism could provide insight into whether and how these models of core-complex formation might be reconciled.

The role of plutonism is significant in the petrotectonic development of many core complexes. Lister and Baldwin (1993) argued that igneous intrusions may trigger transient metamorphism and ductile deformation in shear zones during continental extension. In the Ruby Mountains–East Humboldt Range core complex, as in the Snake Range and Kern Mountains, Nevada (Gottlieb et al., 2022), Cretaceous and Tertiary plutonic rocks can provide fundamental information on the composition and age of the Precambrian basement, the role of mantle-derived magmas during magma petrogenesis, deep-crustal magmatic differentiation processes (fractional crystallization, assimilation, and partial melting), and the relationship between tectonic regime (contraction and extension) and magmatism in the hinterland of the Sevier orogenic belt. Our integrated field, structural, geochronologic, geochemical, and isotopic study of a group of middle to late Eocene and early Oligocene plutonic rocks exposed in the infrastructure of the Ruby Mountains–East Humboldt Range core complex helps address these relationships (Fig. 2).

These intrusive igneous rocks provide an excellent perspective on the transition from thickened continental crust in the Late Cretaceous to widespread continental magmatism in the middle Cenozoic to crustal extension and core-complex formation in the late Cenozoic (early Oligocene–middle Miocene). The Paleogene magmatism that we describe exemplifies the relationship between Eocene–Oligocene magmatism and core-complex formation, a relationship that is complex, involving multiple magmatic phases, a significant time interval (Paleocene–Miocene), extensive partial melting, widespread metamorphism, a triclinic flow pattern, and multi-phase progressive plastic-to-brittle extensional deformation.

The Ruby Mountains–East Humboldt Range metamorphic core complex is a window into thickened continental crust that includes the “root zone” of the Sevier fold-and- thrust belt (Miller and Gans, 1989; Allmendinger, 1992). Folded pre-metamorphic thrust faults (Howard, 1966, 1987; McGrew, 2018), large recumbent fold nappes (Howard, 1980; McGrew and Snoke, 2015; McGrew, 2018), and moderate-pressure to Barrovian higher-pressure metamorphism (e.g., Hodges et al., 1992; Lewis et al., 1999; McGrew et al., 2000; Cooper et al., 2010; Hallett and Spear, 2014) are manifestations of Late Cretaceous crustal thickening in the Sevier hinterland (Coney and Harms, 1984).

Crustal thickening in the hinterland of the Sevier fold-and-thrust belt is thought to have resulted in a high “Nevadaplano” plateau comparable to the Altiplano-Puna plateau of the Andes (DeCelles, 2004; Whitney et al., 2004a; Best et al., 2009; Mix et al., 2011; Cassel et al., 2014). Various studies based on geologic, paleontological, geochemical, and stable-isotopic data (e.g., Wolfe et al., 1998; Dilek and Moores, 1999; DeCelles, 2004; Best et al., 2009; Sewall and Fricke, 2013; Snell et al., 2014; Cassel et al., 2014; Chapman et al., 2015) support the concept that a thickened crust and orogenic plateau existed in the present-day Great Basin from Late Cretaceous until at least late Oligocene time (Colgan and Henry, 2009; Henry et al., 2012; Cassel et al., 2014).

During the middle Tertiary, predominantly felsic magmatism swept southward from British Columbia, Canada, to southern Nevada and an analogous magmatic zone migrated northward from northern Mexico and southern Arizona, USA (Christiansen and Yeats, 1992; Humphreys, 1995; Dickinson, 2006; Fig. 1). The “ignimbrite flare-up” (Coney, 1978) included large ash-flow eruptions from calderas in central Nevada and western Utah (Best et al., 2009; Henry et al., 2012). Some of the ash flows channeled into paleovalleys, which distributed the tuffs across and beyond the plateau, spread to the west across the Sierra Nevada and to the east across eastern Nevada (Henry, 2008; Henry et al., 2012). Colgan and Henry (2009) proposed that the orogenic plateau collapsed in the middle Miocene (ca. 17–16 Ma). This collapse coincided with the eruption of flood basalts of the Steens Mountain Volcanics and Columbia River Basalt Group, incursion of the Yellowstone hot spot, and the final delamination of the shallow-dipping Farallon slab from beneath the western United States of America. Also, this was a time of major plate-boundary reorganization from subduction to transtensional strike-slip faulting along the western margin of North America (Christiansen and Yeats, 1992). Extension associated with the development of the modern Basin and Range province occurred after ca. 10 Ma and has continued to the present (Colgan and Henry, 2009).

The Ruby Mountains are a NNE-striking horst in the late Cenozoic Basin and Range province of northeastern Nevada (Fig. 2). The Ruby Mountains and adjoining East Humboldt Range form one of the larger and most deeply exhumed North American Cordilleran metamorphic core complexes (Crittenden et al., 1980; Armstrong, 1982; Fig. 1). The ranges display the classic core-complex subdivisions: (1) a core of non-mylonitic metamorphic and igneous rocks (infrastructure); (2) a hundreds-of-meters-thick mylonitic shear zone that attenuated the crustal section and forms a carapace structurally above the infrastructure; and (3) an upper tier of locally metamorphosed to non-metamorphosed rocks separated from the metamorphic and igneous terrane by a plastic-to-brittle, normal-sense detachment-fault system.

The infrastructure of the core complex exposes a mid-crustal terrane of Neoproterozoic to Mississippian miogeoclinal strata that were deeply buried, penetratively deformed, metamorphosed, and intruded by granitic rocks prior to exhumation in Miocene time (Howard, 1980, 2000, 2003; Snoke and Lush, 1984; Snoke et al., 1997, 2022; Howard and MacCready, 2004; Colgan et al., 2010; Howard et al., 2011; McGrew and Snoke, 2015; McGrew, 2018). Late Cretaceous and smaller volumes of middle Eocene to early Oligocene granitoids form ~50% of the complex and as much as 90% or more of the deepest exposed levels (Howard, 1966, 1980, 2000; Howard et al., 1979, 2011; Wright and Snoke, 1993; MacCready et al., 1997; Howard and MacCready, 2004). The dominant volume of Late Cretaceous granites is two-mica, coarse-grained to pegmatitic leucogranite that varies from gneiss intimately interlayered and folded with metasedimentary rocks to cross-cutting dikes and sheets (Howard, 1966, 1980, 2000; Lee et al., 2003, their KPG unit; Howard et al., 2011). The infrastructure reached upper amphibolite-facies metamorphic conditions in the Late Cretaceous as dated by tectonic foliation in granite orthogneiss (ca. 92 Ma) that is cut by a gneissic pegmatitic leucogranite (ca. 83 Ma) (Lee et al., 2003; Howard et al., 2011). The mineral assemblage in pelitic schist includes muscovite, alkali feldspar, biotite, sillimanite, garnet, and plagioclase. Assemblages in calc-silicate paragneiss include calcite, plagioclase, alkali feldspar, diopside, tremolite, biotite, and scapolite.

Renewed magmatic activity in Eocene time involved the emplacement of gabbro through leucocratic monzogranite magmas and lasted until the early Oligocene (ca. 29 Ma) when the last sparse but widespread monzogranitic sheets and dikes were emplaced (Wright and Snoke, 1993; McGrew et al., 2000; Howard et al., 2011; McGrew and Snoke, 2015; McGrew, 2018; Sicard and Snoke, 2020; Snoke et al., 2021, 2022; Zuza et al., 2021, 2022). The plutonic rocks and their hosts are mylonitized in the westward-dipping mylonitic shear zone that crops out along the western flank of the Ruby Mountains and East Humboldt Range (Howard, 1980, 2000; Snoke et al., 1997, 2021, 2022; Sicard and Snoke, 2020). The metasedimentary stratigraphic section is thinned to as little as 5% of original thickness in the shear zone (Howard, 1966, 2000). Granitic and metasedimentary rocks in the shear zone exhibit spectacular mylonitic textures, and kinematic indicators consistently indicate top-to-the-WNW shear and stretching (Snoke and Lush, 1984; Lister and Snoke, 1984; MacCready, 1996; McGrew and Casey, 1998). A low-angle, brittle detachment fault zone is superposed on the mylonitic shear zone and thus is a later extensional feature in the exhumation of the core complex (Snoke et al., 2021; Zuza and Dee, 2023).

Mesozoic plutonism in the core complex began with intrusion of small Late Jurassic plutons. One is a granite (ca. 153 Ma) in the southern Ruby Mountains; another is a granodiorite gneiss (ca. 161 Ma) in the lower Lamoille Canyon area (Hudec and Wright, 1990; Hudec, 1992; Howard et al., 2011). After a long hiatus, voluminous Late Cretaceous peraluminous granite was emplaced in an intricate web of sheets, dikes, and other small intrusions. The Late Cretaceous granites make up close to half of the rock exposed in the core complex. The oldest of these rocks is equigranular biotite granite orthogneiss in upper and central Lamoille Canyon, dated by U-Pb zircon and monazite to ca. 91 Ma (Howard et al., 2011; Hetherington et al., 2017; this report), and, in lower Lamoille Canyon, the sillimanite-two-mica granite gneiss of Thorpe Creek (ca. 91 Ma) (Howard, 2000; this report). The equigranular biotite granite orthogneiss in Lamoille Canyon was interpreted as derived by biotite-dehydration melting of graywacke-like gneiss compositions (Lee et al., 2003). In upper Lamoille Canyon, these granites (ca. 91 Ma) are intruded by gneissic banded to pegmatitic leucogranite, which forms the most voluminous igneous rock in the core complex (Lee et al., 2003; Howard et al., 2011; this report). The pegmatitic leucogranite contains biotite, muscovite, and commonly sillimanite or garnet. Single-grain U-Pb dating of these gneissic rocks form spreads of zircon ages from ca. 76 to 69 Ma (Howard et al., 2011; this report). Based on geochemical modeling, muscovite-dehydration melting of metapelite, at <35 km depth, was the process responsible for the petrogenesis of these Late Cretaceous peraluminous granites (Lee et al., 2003). No mantle mass or heat contribution was deemed necessary to produce these leucogranites according to Lee et al. (2003), who interpreted the origin of this granite as related to Late Cretaceous crustal thickening and anatexis of underlying Neoproterozoic metapelitic rocks. Cretaceous crustal anatexis may have been a response to orogenic thickening in the hinterland of the Sevier fold-and-thrust belt (Miller and Gans, 1989; Patiño Douce et al., 1990; McGrew, 1992; Lee et al., 2003) or to subsequent crustal collapse and decompression (Hodges and Walker, 1992; Camilleri and Chamberlain, 1997; Vanderhaeghe et al., 1999; Whitney et al., 2004a; Wells and Hoisch, 2008; Chapman et al., 2021). The common presence of middle Paleogene zircon rims, together with geochemistry of the dated zircon, indicated to Howard et al. (2011) that the Cretaceous pegmatitic leucogranite likely underwent partial anatexis, some remobilization, and recrystallization when reheated and affected by fluids in middle Paleogene time.

Eocene to early Oligocene magmatism lasted from ca. 40 to 29 Ma (Wright and Snoke, 1993; Premo et al., 2005, 2008, 2014; Howard et al., 2011; Romanoski, 2012; Moscati et al., 2023; this report). The largest of the Eocene intrusions is the Harrison Pass pluton (ca. 36 Ma), which forms an ~140 km2, composite pluton in the southern Ruby Mountains (Burton, 1997; Barnes et al., 2001; Colgan et al., 2010; Fig. 2). This compositionally diverse pluton exposes an early suite of biotite ± hornblende granodiorite, biotite monzogranite including porphyritic roof dikes, and a late-stage monzogranite of Green Mountain Creek locally characterized by two micas and by restitic, surmicaceous enclaves (Burton, 1997; Barnes et al., 2001). Mafic to intermediate-composition rocks constitute an important part of the pluton, both as early syn-plutonic dikes and mafic enclaves and as dikes emplaced late in the magmatic history of the pluton. This epizonal to mesozonal pluton originated when emplacement of mafic magmas into the lower crust caused early crustal melting and hybridization with mafic magma, followed by later crustal melting of pelitic and semipelitic source rocks that produced two-mica granites (Barnes et al., 2001).

Elsewhere in the core complex, middle to late Eocene intrusive bodies are smaller but are widespread (e.g., Howard, 1966; McGrew, 1992, 2018; Wright and Snoke, 1993; Sicard and Snoke, 2020; Zuza et al., 2021; Snoke et al., 2022). They encompass a wide range of bulk compositions, similar to the Harrison Pass pluton. Mafic to intermediate-composition rocks include scant hornblende-pyroxene gabbro, biotite-hornblende quartz diorite, and hornblende-biotite tonalite to granodiorite. Silicic rocks range from hornblende-biotite monzogranite to garnet two-mica leucomonzogranite and muscovite pegmatite. Coeval middle Eocene mafic volcanic and volcaniclastic rocks crop out in the hanging wall of the Ruby Mountains–East Humboldt Range detachment fault in the southeastern East Humboldt Range (Taylor, 1984; Brooks et al., 1995a, 1995b).

Distinctive biotite ± hornblende monzogranite forms sheets and dikes. Where dated by U-Pb (zircon) from four sites in the core complex, this rock type showed a narrow age range ca. 32–29 Ma (Wright and Snoke, 1993; MacCready et al., 1997; Zuza et al., 2021; Snoke et al., 2022). Our new U-Pb dating expands its age range to ca. 40 Ma.

Whether or how the Paleogene magmatism is related to tectonism, and particularly to unroofing of the core complex, has been disputed. Although the timing of initiation of extension and unroofing of the Ruby Mountains–East Humboldt Range core complex remains uncertain, McGrew and Snee (1994, their figure 7) suggested that exhumation may have begun in the late middle Eocene after a heating event at ca. 40 Ma, manifested by the intrusion of mafic magmas. Biotite K-Ar and 40Ar/39Ar age “chrontours” young westward from 35 to 20 Ma across the Ruby Mountains–East Humboldt Range core complex (Kistler et al., 1981) and have been interpreted as recording progressive westward exhumation during Eocene to early Miocene extensional faulting (McGrew and Snee, 1994; Gifford, 2008). An alternative interpretation proposes that Eocene magmatism and thermal softening led to Oligocene diapiric upwelling with a general-shear mylonitic carapace (Zuza et al., 2022; Zuza and Cao, 2023; Levy et al., 2023).

5.1 Introduction

Compositionally diverse middle Paleogene plutonic rocks in the upper Lamoille Canyon area of the Ruby Mountains encompass gabbro, tonalite, biotite monzogranite, two-mica leucomonzogranite, and pegmatitic leucocratic syenogranite. Together they make up perhaps 5%–10% of the exposed infrastructure. Medium-grained biotite monzogranite is the most prominent and voluminous lithotype, typically with <10% mafic minerals and pegmatite is common. Geochemical and isotopic data presented in section 8 indicate that two categories of Paleogene granitic rocks are present, although recognition is generally not possible at the outcrop scale. The most abundant of the two types is here referred to as monzogranite of Snow Lake Peak type, after prominent outcrops in upper Lamoille Canyon, Island Lake cirque, Castle Lake, and upper Thomas Canyon (Fig. 3). Other intrusions of Snow Lake Peak–type biotite monzogranite form scattered dikes and sheets a few meters to tens of meters thick in other areas of the Ruby Mountains–East Humboldt Range metamorphic core complex (Figs. 4A4C; Howard, 2000; Howard and MacCready, 2004; McGrew and Snoke, 2015; McGrew, 2018; Sicard and Snoke, 2020; Zuza et al., 2021).

The other type of granite is referred to as the Overlook type, after a road-cut exposure near the Glacier Overlook site on the Lamoille Canyon road. Most granites of this type are leucocratic monzogranite, but they range to leucocratic syenogranite. Many display pegmatite or aplitic features, and they lack the flaggy weathering characteristic of Snow Lake Peak–type monzogranites, although they intrude and are intruded by the Snow Lake Peak–type monzogranite. Unlike Snow Lake Peak–type intrusions, Overlook-type granites occur as dikes and as leucosome-like segregations in Late Cretaceous pegmatitic gneiss (Fig. 5). Many of the dikes are undeformed, although aplite dikes are in some cases foliated parallel to their walls. Muscovite, biotite, and, rarely, garnet occur in varying but generally low amounts, and some dikes lack any mafic phases. Most of these dikes are thinner than 1 m, with the following exception. A common and distinctive group of very coarse-grained pegmatite dikes exhibit randomly oriented sheaves of muscovite or biotite which give an overall “chicken-track” appearance to the rock (Fig. 4D). The chicken-track pegmatite dikes commonly exceed 1 m in thickness and rarely attain 10 m. Where a swarm of chicken-track pegmatite dikes crosses one 45-m-wide outcrop, they constitute 26% of the outcrop (Lee and Barnes, 1997).

In upper Lamoille Canyon, the pegmatite and aplite dikes tend to cluster in three nearly orthogonal orientation groups: (1) dominant steep and east striking, (2) subhorizontal, and (3) subordinate steep and north striking (Figs. 6A and 6B). Small scarce quartz veins also tend to fall in these orientations. The dikes in the subhorizontal group dip at low angles to the east, roughly parallel to foliation in the marble and granite gneiss country rocks. Among the dominant steep, east-striking dike orientations, chicken-track pegmatite dikes commonly strike slightly north of east, whereas other, mostly older steep dikes of the suite tend to strike east or slightly south of east. This pattern suggests a progressive slight counterclockwise shift in azimuth of east-west dike emplacement.

5.2 Mapped Intrusions of Snow Lake Peak Type

Our mapping focused on well-exposed areas of the infrastructure in the upper Lamoille Canyon and Castle Lake areas (Fig. 3). Two Snow Lake Peak–type biotite monzogranite intrusions mapped here as the Snow Lake Peak (Fig. 7) and Castle Lake (Fig. 8) intrusions exemplify the intrusive style, some of the compositional variation, and deformation of this notable lithotype. Orientations and compositions of numerous smaller dikes and irregular bodies of the intrusive sequence that crop out within these study areas and elsewhere in Lamoille Canyon further enrich the understanding of the middle Paleogene intrusive episode.

The Snow Lake Peak and Castle Lake intrusions intrude metamorphosed marble and calc-silicate rocks (Cambrian and Ordovician marble of Verdi Peak), older but structurally overlying feldspathic, micaceous metaquartzite (metamorphosed Neoproterozoic and Cambrian Prospect Mountain Quartzite), and the network of Late Cretaceous granitic orthogneiss and gneissic pegmatitic leucogranite (Fig. 9). Foliation and layering in all these country rocks dip gently regionally and exhibit some mesoscopic fold hinges.

5.2.1 Snow Lake Peak Intrusion

The Snow Lake Peak intrusion forms an east-dipping sheet consisting chiefly of medium-grained biotite monzogranite with subordinate two-mica leucomonzogranite. Contacts with the adjacent wall rocks are sharp. The thickness of the sheet varies from ~90 to 40 m (Fig. 7C). Many biotite monzogranite dikes, related to the sheet, intrude country rocks that roof the sheet. Such dikes vary from massive and non-foliated to ones with foliation subparallel to the dike walls, and they commonly truncate older foliation or compositional layering in the country rocks. The orientation of the dikes varies from steeply dipping to a distinct group of dikes that have low to moderate dips (Figs. 6C and 6D).

Two-mica leucomonzogranites are a prominent part of the Snow Lake Peak intrusion. Although muscovite is readily identified in hand sample, it is less abundant than biotite. Garnet is a common accessory phase. These rocks crop out as dikes cutting the metamorphic wall rocks as well as the Snow Lake Peak intrusion. Thin pegmatite selvages occur along the margins of many leucomonzogranite dikes (Figs. 4A4C), and some leucogranites grade into pegmatite. Along Lamoille Creek in uppermost Lamoille Canyon, the Snow Lake Peak intrusion is capped by a distinctive shallowly dipping, two-mica leucomonzogranite sheet, as much as 20 m thick, that is characterized by 1-cm-diameter garnet. The capping leucomonzogranite is in sharp intrusive contact against older biotite monzogranite of the main intrusion. Dikes of garnetiferous, two-mica leucomonzogranite emanate from the cap and cross-cut wall-rock compositional layering and foliation at a high angle.

The Snow Lake Peak intrusion hosts small enclaves of metasedimentary rocks, some of them portrayed on the geologic map (“en” in Fig. 7A). Intrusion breccia is developed at one locality along the lower contact of the intrusion, where the biotite monzogranite forms an intricate network into the host Cretaceous pegmatitic leucogranite orthogneiss, resulting in irregularly shaped, commonly angular orthogneiss xenoliths ranging from a few centimeters to >1 m in size. This intrusive breccia indicates that host rocks fractured brittlely during emplacement of the biotite monzogranite.

Schlieren layering (Fig. 10A) occurs at scattered sites in the biotite monzogranite of the Snow Lake Peak intrusion, particularly near the roof and floor of the intrusion. The schlieren consist of alternating centimetric-thick layers of biotite-rich and biotite-poor monzogranite. Biotite is generally oriented parallel to layering. Schlieren layers locally transect each other in shapes similar to cross bedding. Schlieren are commonly interpreted as a result of flow segregation during viscous flowage of a partially crystallized magma (e.g., Vernon and Paterson, 2008; Ardill et al., 2020; Clemens et al., 2020).

A distinctive group of rare hornblende leucomonzogranite bodies crop out in upper Lamoille and Thomas Canyons. They are distinguished by centimeter-scale clusters of hornblende ± biotite in a medium-grained biotite-bearing leucomonzogranite matrix, which give the rocks a spotted appearance (Fig. 10B). The dark spots are commonly rimmed by aureoles that lack mafic minerals. These hornblende-spotted rocks occur in leucomonzogranite dikes that intrude marble or calc-silicate host rocks, which suggests the possibility that amphibole stability was enhanced by minor local assimilation of Ca-rich material. If this is the case, then amphibole was stabilized in the parental magmas in situ; amphibole was not crystallized at deeper levels and carried to the level of emplacement as phenocrysts. This relationship has important implications for barometric measurements.

The biotite monzogranite is medium grained, equigranular, and massive to lineated and/or foliated. Mafic content typically is 5%–8%, extending to <5% in leucomonzogranite variants. Typical biotite monzogranite consists of biotite, plagioclase, alkali feldspar, and quartz ± calcic amphibole, with minor amounts of Fe-Ti oxides and various accessory minerals (chiefly apatite, allanite, epidote, monazite, and zircon) (Figs. 11A11D). Plagioclase is normally zoned from An35 to An24 cores and An23 to An18 rims, and the amphibole ranges from ferroedenite to hastingsite, with Mg/(Mg + Fe) from 0.22 to 0.29 (File S1 in the Supplemental Material1). Muscovite occurs as a late-stage alteration mineral, and chlorite locally replaces biotite.

The leucomonzogranites contain primary muscovite and biotite, local garnet, and, in rocks that intrude marble, clusters of calcic amphibole (described in two paragraphs above in section 5.2.1; Figs. 10B and 11B). Plagioclase in the leucomonzogranite has core compositions from An28 to An17 and rim compositions that average An17. Myrmekitic intergrowths, associated with alkali feldspar, are common.

5.2.2 Castle Lake Intrusion

The Castle Lake intrusion crops out ~4 km south of the Snow Lake Peak intrusion (Fig. 3). This southern biotite monzogranite intrusion forms a curviplanar sheet-like shape (Strike, 2000) of ~1 km2, exposed through a relief of ~244 m (Figs. 8A8C). This intrusion is locally discordant to foliation and layering in the adjacent host metasedimentary rocks and Cretaceous and Tertiary granitic rocks.

The intrusion contains a prominent zone rich in country-rock xenoliths, from centimeters to meters in width, and local auto-intrusion breccia rich in enclaves of early phases of the monzogranite (Strike, 2000). The xenolith-bearing zone lies beneath xenolith-poor parts of the intrusion and is ~120 m across and 20 m thick. It forms a stratigraphic-like sequence of three subzones defined by variation of xenolithic rock type. The lowest subzone is rich in xenoliths of calc-silicate rock and leucogranite orthogneiss. Small sillimanite-biotite schist xenoliths are concentrated in the middle subzone. Xenoliths in the highest subzone consist of feldspathic micaceous metaquartzite and leucogranite orthogneiss. This lithologic sequence of the zones broadly mimics the structural order of the country rocks, but the metaquartzite and sillimanite-biotite schist xenoliths occur below their country-rock counterparts. Furthermore, sillimanite-biotite schist in the country-rock sequence is only a few meters thick and present only locally, in contrast to much more abundant, country-rock lithotypes of marble, granitic gneiss, calc-silicate rock, and micaceous quartzite. Foliation in the xenoliths varies in orientation from one xenolith to the next, indicating that the fragments were rotated during emplacement of the enclosing biotite monzogranite. There is little evidence that the xenoliths were partially assimilated into the magma given that they are angular and lack reaction rinds. We interpret the xenoliths as evidence of magmatic stoping where roof or wall-rock blocks were incorporated into the monzogranite magma. When considered along with the intrusion breccia locally developed along the floor of the Snow Lake Peak intrusion, these observations suggest that brittle fracturing played a role in the emplacement of the biotite monzogranite magma.

5.3 Leucocratic Granite, Pegmatite, and Aplite of Overlook Type

Overlook-type granites crop out as dikes that cross-cut older rocks of the infrastructure and as leucosome-like bodies primarily associated with Late Cretaceous granitic gneisses. Some of the leucosome-like bodies cross-cut foliation in the host gneiss (Figs. 5A and 5B), whereas others are parallel to or interfinger with foliation (Fig. 5C). The typical Overlook type is leucocratic monzogranite with less-common syenogranite, is medium grained, equigranular, and weakly to non-foliated, and has a low color index (mostly <5). Aplitic and pegmatitic textures are less common, with grain sizes from ~1 mm to >20 cm; some leucomonzogranite grades into pegmatite. Biotite and muscovite are conspicuous in the medium-grained Overlook type, and Mn-rich garnet is locally an accessory mineral. Many pegmatites are garnet bearing with more muscovite than biotite.

5.4 Tonalitic Rocks

Dark, fine-grained, biotite-rich tonalitic rocks are scattered throughout the upper Lamoille Canyon area. They are herein referred to as “tonalitic rocks” although they vary to quartz diorite. The tonalitic bodies are generally <2 m thick. Their extremely irregular, non-planar, anastomosing outlines contrast with the tabular shapes of most granitic dikes of the intrusive sequence. Some of the tonalitic rocks exhibit schistose foliation, which is folded locally into mesoscopic isoclinal folds. The tonalitic rocks are not voluminous, occupying no more than ~1% of the area shown in Figure 7A.

The tonalitic rocks show both younger and older intrusive relationships relative to the biotite monzogranite of the Snow Lake Peak intrusion, indicating that they were emplaced during the same interval. Locally, the tonalitic rocks are cut by dikes of garnet-two-mica monzogranite or garnet-bearing, two-mica leucogranite, which may enclose xenoliths of the foliated, dark tonalitic dike rock. Elsewhere, the dark rocks cut muscovite-bearing biotite monzogranite. Mingling relationships between the dark tonalitic rocks and monzogranite of Snow Lake Peak type are common (Figs. 10C and 10D).

The dark, fine-grained, biotite-rich tonalitic rocks vary from tonalite to quartz diorite in terms of modal classification. In a geochemical sense, these rocks are much more potassic than tonalite or quartz diorite owing to the large proportions of biotite (28–36 vol%). Amphibole (0–6 vol%) is also common (Figs. 11B and 11D) and is predominantly ferrotschermakite and ferroedenite, with values of Mg/(Mg + Fe) from 0.40 to 0.50. Plagioclase ranges in composition from An53 to An18. Both normal and reverse zoning were observed (File S1 [see footnote 1]). Accessory minerals are apatite, titanite, Fe-Ti oxides, and zircon.

5.5 Gabbroic Rocks

Clustered small pods and sheets of amphibole ± pyroxene gabbro a few meters thick intrude the marble of Verdi Peak and Cretaceous granitic gneisses in some areas of the northern Ruby Mountains (Howard, 1966, 2000; Smith and Howard, 1977; Howard and MacCready, 2004; Snoke et al., 2022). They are scarce in the areas of our detailed maps (Figs. 7 and 8), with a few outcrops west of Liberty Pass, but concentrate ~6 km NNE of Snow Lake Peak. The gabbros studied here occur in the latter area, where they are locally intruded by gneissic pegmatitic leucogranite and are surrounded by the folded Prospect Mountain Quartzite at the nose of the Lamoille Canyon fold nappe (Fig. 9).

The gabbroic rocks have color indices >50. They generally lack penetrative deformation fabric. The rocks are typically medium to coarse grained; scarce pegmatitic varieties also occur. Their compositions include hornblende gabbronorite and hornblende-orthopyroxene-augite quartz gabbro, with or without biotite. In some samples, hornblende is clearly magmatic, whereas in others it is deuteric after pyroxene. Cummingtonite, anthophyllite, and(or) phlogopite may be present. Plagioclase varies from andesine to labradorite. Accessory minerals include alkali feldspar, titanite, and magnetite. The rocks are commonly altered to widespread chlorite.

6.1 Foliation and Lineation

Foliation or compositional layering is nearly pervasive in the metasedimentary host rocks (Fig. 12A). The foliation is penetrative, whereas compositional layering is manifested by mineralogical segregation. The Cretaceous orthogneisses also exhibit penetrative foliation or mineralogical segregation layering.

The biotite monzogranite (Snow Lake Peak type), although massive in outcrop, is commonly foliated and lineated. Only locally is the biotite monzogranite strongly foliated. The foliation in the biotite monzogranite is typically defined by platy biotite grains and to a lesser extent by flattened quartz or feldspar grains (Figs. 11A, 11C, and 12B). At one locality, the foliation is isoclinally folded. Foliation locally transects compositional boundaries, indicating that the fabric in the biotite monzogranite developed chiefly in the solid state, with little or no melt present. The biotite monzogranite thus deformed as a coherent polymineralic mass with quartz being the weakest phase and feldspar the strongest.

In some samples, the lineation, defined chiefly by elongated quartz and biotite aggregates, is considerably more pronounced than foliation, yielding a L > S tectonite. The L-S fabric is related chiefly to solid-state deformation given that it crosses local compositional boundaries in the biotite monzogranite. However, foliation is parallel to dike walls in some aplite and biotite monzogranite dikes. In monzogranite that contains hornblende ± biotite aggregates (Figs. 10B, 11B, and 11D), the hornblende-rich spots are commonly elongated giving the rock a lineation. This lineation is interpreted as magmatic rather than solid state based on lack of deformation of the hornblende.

Distinctive mineral lineations, typical of the core complex infrastructure, include aligned sillimanite in the Late Cretaceous pegmatitic orthogneiss and in feldspathic micaceous metaquartzite of the Prospect Mountain Quartzite; lineations typically trend north-south (Fig. 12C). The mineral lineations typically parallel fold hinges that are at least in part Cretaceous (Howard, 1980; Howard et al., 2011), yet MacCready et al. (1997) reported the lineation orientation to be present also in an Oligocene monzogranite (ca. 29 Ma) in the Soldier Peak quadrangle (Fig. 2; Snoke et al., 2022). Elongation lineations in Eocene biotite monzogranite in the Snow Lake Peak area instead mostly trend ESE (Fig. 12C), which is colinear to the mylonitic lineation in the overlying 200–300-m thick mylonitic shear zone (Howard, 2000).

6.2 Mylonitic and S-C Fabrics

Although mylonitic foliation and lineation are best developed and pervasive in the mylonitic shear zone that forms a carapace on the infrastructure, scattered mylonitic zones occur within the infrastructure. These mylonitic fabrics occur in the typical biotite monzogranite as well as garnet-bearing, two-mica leucocratic monzogranite.

These zones of local mylonitization in the infrastructural core have lineations oriented WNW-ESE, colinear with the elongation lineations of the mylonitic lineation of the overlying carapace shear zone, and are interpreted to be small-scale shear zones that are kinematically related to the carapace shear zone. However, the sense of shear in the infrastructural mylonitic rocks is variable, either to the WNW or ESE. S-C fabrics are locally well developed in biotite monzogranite, and their presence there, but not in the immediate host rocks, suggests to us that biotite monzogranite bodies served as strain guides in the infrastructural core during late mylonitization and exhumation of the core complex. The C-surface is defined by grain-size-reduced biotite and quartz with the biotite smeared out along the C-surface plane. The local development of S-C fabric in dikes of the biotite monzogranite may be syn-magmatic given that an undeformed pegmatite dike cuts the fabric at one key locality. At this locality, non-penetrative slickenlines on the cross-cutting pegmatite dike are colinear with the penetrative lineation on C-surfaces of the composite S-C fabric in the adjacent biotite monzogranite. These relationships indicate a structural coupling between the overlying mylonitic shear zone and subjacent infrastructure.

7.1 Introduction

Published ages from the Ruby Mountains and Lamoille Canyon area allowed workers to document Late Cretaceous magmatism followed by Paleogene magmatism, with a potential subdivision of Paleogene magmatism into distinct Eocene and Oligocene events (Wright and Snoke, 1993; Sullivan and Snoke, 2007). More-recent U-Pb dating of zircon by sensitive high-resolution ion microprobe–reverse geometry (SHRIMP-RG) mass spectrometry (Stanford University–U.S. Geological Survey [USGS]) (Howard et al., 2011) identified significant intra-grain complexity, which included the possibility of overlap of the Eocene and Oligocene events and of Paleocene magmatism at ca. 60 Ma. To clarify the relationships between magmatic, tectonic, and thermal events, additional analytical work was conducted by SHRIMP-RG, laser-ablation multi-collector inductively coupled plasma mass spectrometry (LA-MC-ICP-MS), and 40Ar/39Ar incremental-release dating.

Samples were processed using standard density and magnetic mineral separation techniques to isolate zircon and monazite. Monazite for LA-MC-ICP-MS and zircon for SHRIMP-RG were handpicked into epoxy mounts, polished, inspected with transmitted and reflected light, and imaged via cathodoluminescence (CL). Zircon analyzed by LA-MC-ICP-MS was subjected to a modified thermal annealing and chemical abrasion procedure (Mattinson, 2005). These treated zircons were mounted and imaged by scanning electron microscope (SEM)–assisted back-scattered electron (BSE) and color CL.

Sample locations are listed in Table 1, and a summary of U-Pb data is given in Table 2. Age and trace-element data along with analytical details for SHRIMP-RG analysis are presented in File S2 (see footnote 1) and in Moscati et al. (2023). Similarly, LA-MC-ICP-MS age and trace-element results on zircon and age data on monazite along with analytical details are presented in File S3. Hafnium isotope results are given in File S4, and File S5 presents supporting data, including CL images, rare-earth-element (REE) plots, and trace-element diagrams.

Th-Pb (monazite) and U-Pb (zircon) isotope ratios and Hf isotope data (zircon) were analyzed by LA-MC-ICP-MS at the University of California, Santa Barbara, using a Photon Machines 193 nm excimer laser and Nu Instruments Plasma HR LA-MC-ICP-MS. For zircon grains, U-Pb analyses were conducted first, using a laser spot diameter of ~14 µm, followed by analysis of Hf isotopes using a 50–55 µm laser spot. The U-Pb ages were used to calculate εHf values at the interpreted age of emplacement. The data were reduced using Iolite software (iolite.xyz), and 91500, Mud Tank, and TEMORA zircon standards were used to monitor instrument performance and calibration. Trace-element abundances were measured by laser-ablation quadrupole inductively coupled plasma mass spectrometry (LA-Q-ICP-MS) at Texas Tech University.

More than 30 samples were collected from Lamoille Canyon and adjacent areas for U-Pb zircon age dating using the SHRIMP-RG. This work was part of USGS-funded projects investigating the geology of the Great Basin and, in particular, Eocene events potentially related to Carlin-type gold deposits (e.g., Hofstra et al., 2010; Moscati et al., 2023). Samples range from gabbro to biotite monzogranite to pegmatitic leucomonzogranitic gneiss. Sample sites are in Lamoille Canyon except one (diorite WRP 02-4) from Secret Creek gorge (~32 km to the NNE; Fig. 2).

The majority of LA-MC-ICP-MS (zircon) results come from the study of Romanoski (2012), which focused on dating multiple samples from individual outcrops displaying clear cross-cutting relationships (File S5). Sampling focused on granitic and metagranitic rocks. Additional LA-MC-ICP-MS data on zircon and monazite are products of undergraduate research at Texas Tech University supervised by Callum Hetherington.

In what follows, age results are presented in approximate age sequence, from oldest to youngest, starting with Late Cretaceous granitic gneisses of the infrastructure, then Paleogene mafic to intermediate rocks, and then Paleogene granites. Among the Paleogene granites, monzogranitic samples are ascribed to Snow Lake Peak or Overlook type in cases where bulk-rock major and trace-element and/or isotope data are available (Table 3). Figures 1318 present weighted average, isochron, and probability plots for these samples. CL images, REE diagrams, and trace-element diagrams may be found in File S5.

7.2 Late Cretaceous and Early Paleogene Granitic Gneisses

Equigranular granitic gneiss was identified by Lee et al. (2003) as the oldest Late Cretaceous granitic gneiss in the Lamoille Canyon area. LA-MC-ICP-MS data on zircon from two equigranular granitic gneiss samples (RM-26 and RD-31) yielded weighted-mean 206Pb/238U (zircon) ages of 91.8 ± 0.4 Ma (mean square weighted deviation [MSWD] = 1.3, N = 20) and 90.0 ± 0.4 Ma (MSWD = 2.3, N = 26), respectively (Figs. 13A and 13B). Data from eight zircons from a third sample (RD-19A) scatter along concordia from ca. 94 to 80 Ma (Fig. 13C). The εHf values of zircon grains dated to 92–90 Ma range from −10 to −26 (File S4).

Sample WRP 02-5 is from the mylonitic granite gneiss of Thorpe Creek (Howard, 2000), a garnet-bearing, leucogranitic gneiss, collected at the mouth of Lamoille Canyon. Zircon dates from oscillatory zoned mantle domains attributed to magmatic crystallization are dominated by ca. 90 Ma dates (Fig. 13D). A few ca. 80 Ma dates and one concordant analysis at ca. 60 Ma from a CL-bright rim were also obtained. Grouping analyses on the basis of textures provides a mean 206Pb/238U crystallization age of 91.2 ± 1.0 Ma (MSWD = 1.6; N = 10; Fig. 13D). The ca. 91 Ma analyses gave variable Th/U ratios, with four analyses giving Th/U values of ~0.01–0.02 and another group of seven yielding values of ~0.4–0.7 (File S2). Chondrite-normalized REE values for the ca. 91 Ma analyses yield a wide range of Eu/Eu* values (where Eu/Eu* expresses the normalized abundance of Eu divided by the expected abundance of Eu when normalized Eu concentration plots on a line connecting normalized Sm and Gd abundances) between 0.1 and 0.9, which negatively correlate with Hf contents, which increase from ~8000 ppm to >20,000 ppm as Eu/Eu* decreases (File S5). Eocene monazite ages were reported from this Thorpe Creek unit by Wright and Snoke (1993), who also recognized complex zircon populations. The ca. 91 Ma zircon age of the granite gneiss of Thorpe Creek overlaps the ca. 92–90 Ma ages determined for equigranular granitic gneiss in upper Lamoille Canyon, suggesting that these units are temporally, and potentially petrologically, equivalent.

A medium-grained gneissic biotite granite collected near Liberty Pass (sample KW-25E) was sampled as a possible equivalent of the equigranular granitic gneiss. However, this sample yielded an LA-MC-ICP-MS weighted-mean 206Pb/238U (zircon) age of 69.87 ± 0.57 Ma (MSWD = 1.8, N = 6; Fig. 14A).

Sample WRP 08-27 is a mylonitic, fine-grained, leucocratic monzogranitic orthogneiss collected from a large outcrop just east of the Powerhouse Picnic Area parking lot near the mouth of Lamoille Canyon (Table 1). The sample was originally thought to be the granite gneiss of Thorpe Creek. However, SHRIMP-RG analysis of zircons yielded two basic age groups: one at ca. 70 Ma and one at ca. 37 Ma (Fig. 14B). The ca. 37 Ma date was originally interpreted as a metamorphic age because these analyses came from a variety of rim zones (some with magmatic zoning) but also are marked by low Th/U values (~0.02; File S2) suggesting crystallization from metamorphic fluids. However, the older ages display significant scatter, from 87.5 to 52.3 Ma (mean 206Pb/238U age = 72.1 ± 4.2 Ma; MSWD = 389; N = 17; Fig. 14B). Thus, these may be inherited grains or magmatic grains affected by Pb loss, or mixed ages. Supporting data show they were obtained mainly from relatively CL-bright, magmatic mantles (File S5) with U contents ~1000 ppm, Th/U ~0.2–0.5, Eu/Eu* ~0.2–0.5, and Hf contents ~11,000 ppm (Files S2 and S5). The younger population yields a mean 206Pb/238U date of 36.99 ± 0.68 Ma (MSWD = 5.6; N = 6; Fig. 14B). Individual analyses are characterized by low Th/U and Eu/Eu* ratios and elevated Hf contents (~20,000 ppm).

Sample site ARRM11-03 is located north of Lamoille Lake (Fig. 7A). At this location, biotite monzogranite (Snow Lake Peak type) intrudes coarse-grained pegmatitic gneiss along a shallowly dipping contact (File S5). Zircon grains from the pegmatitic gneiss (sample ARRM11–03c) are of two types: ones with mottled cores and thin, CL-bright oscillatory-zoned rims and prismatic grains with weakly zoned, CL-dark cores and weakly zoned, CL-bright rims (File S5). Analysis of zircons with mottled cores and one prismatic grain with CL-dark core by LA-MC-ICP-MS yielded three age populations at 71.1 ± 1.7, 52.3 ± 3.5, and 37.9 ± 3.6 Ma (Table 2; File S3, Fig. 14C). The oldest ages are from the core zones of an elongate prism, intermediate ages are from or adjacent to mottled cores, and the youngest ages are all from crystal rims (File S5). The other prismatic grains yielded a range of Paleozoic ages from 519 to 471 Ma (File S3). Cores of some grains with mottled cores yielded εHf values <−42, whereas the ca. 71 Ma grains yielded values of ~−24.5, and a Paleozoic grain yielded a single εHf value of −12.7.

Sample site ARRM11-11 is located on the floor of upper Lamoille Canyon ~500 m southwest of the trailhead parking area (Fig. 7A), where a very coarse-grained pegmatitic gneiss intrudes and cuts foliation in an older banded gneiss (File S5). LA-MC-ICP-MS analysis of mainly elongate zircon prisms from the banded gneiss (sample ARRM11-11a) yielded 39 concordant dates with a weighted-mean value of 71.10 ± 0.36 Ma and MSWD = 3.1 (Table 3; Fig. 15A). Two grains yielded younger, slightly discordant dates (Fig. 15A). Hf abundances are ~12,000 ppm and Eu/Eu* is ~0.3. The εHf values of these grains range from −28.4 to −36.3 (average −31.6 ± 0.76).

Zircon from the pegmatitic gneiss (sample ARRM11-11b) display varied CL zoning patterns and mottling (File S5); LA-MC-ICP-MS analysis of 19 grains yielded a weighted-mean age of 60.64 ± 0.20 Ma with MSWD = 0.48. The population also contained a discordant, probably antecrystic grain and two concordant zones with ages younger than 35 Ma (Fig. 15B). The ca. 60 Ma age result is of interest because it is similar to rare ca. 60 Ma ages reported by Howard et al. (2011) and supports the idea that a distinct magmatic event occurred at that time. Hf ranges from 11,000 to 14,000 ppm and Eu/Eu* varies from ~0.1 to 0.3. The εHf values of these grains range from −36.1 to −52.7 (average −46.3).

Sample site ARRM11-14 is located near the east fork of Lamoille Creek adjacent to the Ruby Crest Trail (Fig. 7A). This outcrop exposes older, foliated, banded granitic gneiss cut by a later, coarse-grained pegmatitic gneiss (File S5). The banded gneiss (sample ARRM11-14a) yielded mainly elongate, prismatic zircon grains, typically with CL-dark interiors and oscillatory-zoned, CL-brighter outer zones (File S5). Thirty-one concordant dates by LA-MC-ICP-MS yielded a weighted-mean age of 70.7 ± 0.53 Ma, MSWD = 5.5 (Table 2; Fig. 15C). Hf abundances cluster at 12,500 ppm and Eu/Eu* is ~0.3. Average εHf was −29.9 ± 2.7. The cross-cutting pegmatitic gneiss (ARRM11-14b) yielded zircons with a range of habits and CL zoning (File S5). Of these, 24 analyses yielded a weighted-mean age of 75.6 ± 1.1 Ma, with a large MSWD of 16 (Table 2), and a subset of these grains gave an average εHf value of −30.6 ± 1.7. Rim analyses of seven slightly discordant grains plot on a chord with lower intercept age of 34.2 ± 2.7 Ma, MSWD = 1.16 (Fig. 15D inset), and an upper intercept age of 419 ± 240 Ma. Hf abundances range from ~12,000 to 15,000 ppm and Eu/Eu* varies from 0.06 to 0.2. The εHf values obtained from these ca. 34 Ma analytical spots is −27.1 ± 2.2. It is noteworthy that the age determined for the cross-cutting pegmatitic gneiss ARRM11-14b is older than that of its host banded gneiss (ARRM11–14a), as illustrated in weighted-average and probability plots (Figs. 15C and 15D). This discrepancy exemplifies the complexity of seemingly magmatic zircon in Ruby Mountains granites.

Sample site ARRM11-16 is located north of the eastern lake of Dollar Lakes (Fig. 7A). The outcrop consists of a banded granitic gneiss intruded by a coarse-grained pegmatitic gneiss, which itself is boudinaged (File S5). Only the pegmatitic gneiss (sample ARRM11-16b) yielded zircon after processing. Most grains display mottled cores and extensive pitting and thin, CL-bright rims (File S5). U-Pb ages (LA-MC-ICP-MS) vary from 71.6 Ma in one grain to 37.2 Ma, with a range of mostly discordant intermediate dates (not shown). Four young, oscillatory-zoned rim analyses yielded a weighted-mean age of 37.9 ± 1.6 Ma, MSWD = 4.5. Sparse trace-element data for this sample consist of two analyses with ~10,500 ppm Hf and Eu/Eu* ~0.35 and three others with >22,000 ppm Hf and Eu/Eu* from 0.06 to 0.16. The average εHf value of these grains is −24.

7.3 Paleogene Mafic and Intermediate Rocks

A quartz diorite (sample WRP 02-4) and two gabbros (samples WRP 02-8 and Rby-103) were dated by SHRIMP-RG. Zircon from these samples is characterized by oscillatory zoning and U and Th contents <1000 ppm, features characteristic of magmatic grains. Most U-Pb analyses were obtained from oscillatory-zoned parts of grains, which permitted calculation of relatively precise ages. Each sample produced a mean 206Pb/238U age between 40 and 36 Ma (Fig. 16): WRP 02-4, 38.77 ± 0.50 Ma, MSWD = 1.07, N = 10; WRP 02-8, 36.61 ± 0.68 Ma, MSWD = 1.6, N = 10; and Rby-103, 38.46 ± 0.27 Ma, MSWD = 3.1, N = 7. The REE patterns of these three samples (File S5) display weak or no Eu anomalies and nearly linear patterns from La to Yb except for positive Ce anomalies, which are also typical of magmatic zircon.

A tonalite from upper Lamoille Canyon, sample H5-LC-110 (Fig. 7A), yielded oscillatory zoned zircon with CL-bright and CL-dark areas and Th/U values 0.1–1.0 (File S5), which indicates crystallization from a melt. The darker-CL zones have high U and Hf contents compared to the brighter zones. SHRIMP-RG analysis yielded a mean 206Pb/238U age of 37.4 ± 1.2 Ma (MSWD = 15; N = 9; Fig. 16D). The large MSWD is attributed to scatter in dates (45–35 Ma) caused by Pb loss from the high-U-Th CL-darker zones (mostly between 1000 and 10,000 ppm; File S5). Chondrite-normalized REE values show a prominent negative Eu anomaly (Eu/Eu* 0.4–0.9) and a distinct concave-downward pattern among the heavy REEs (File S5).

7.4 Paleogene Monzogranitic Rocks of Snow Lake Peak Type

Zircon from two biotite monzogranites of Snow Lake Peak type were analyzed by SHRIMP-RG. The first was collected adjacent to the road in upper Lamoille Canyon (sample WRP 08-26; Table 1) and the other (H6-Rby-101) from upper Lamoille Canyon near the outcrop of the Snow Lake Peak intrusion (Fig. 7A). Eighteen analyses from mainly clear, oscillatory zoned domains of zircon in sample WRP 08-26 yielded a mean 206Pb/238U age of 36.68 ± 0.95 Ma (Fig. 17A). The dates are dispersed between 42 and 30 Ma and yield an MSWD of 81. Trace-element analyses show a wide spread of U and Th abundances, and the spread of dates is attributed to Pb loss (File S2). Omitting the four most-discordant analyses gives a similar date of 36.23 ± 0.53 Ma with a lower MSWD of 20. The Th/U values are 0.1–1.0, and chondrite-normalized REE abundances display negative Eu anomalies (Eu/Eu* 0.08–0.2; File S5) and concave-downward patterns similar to those of tonalite H5-LC-110 (File S5). Sample H6-Rby-101 yielded 30 analyses from 24 zircon grains. Seven analyses are from magmatic, mainly CL-bright mantle zones and CL-dark euhedral rims. Dates range between 45 and 38 Ma (Fig. 17B) with five analyses used to define a mean 206Pb/238U age of 39.9 ± 1.4 Ma (MSWD = 3.5; N = 5). The U abundances are <1000 ppm and Th is <500 ppm, yielding Th/U values of ~0.5 that are compatible with typical magmatic zircon (File S5). The analyzed domains have low Hf contents and variable Eu/Eu* values with concave-downward REE patterns (File S5).

Zircon grains from biotite monzogranite ARRM11-03b1 display a range of CL features and zoning patterns (File S5). Nevertheless, 17 LA-ICP-MC-MS analyses resulted in a weighted-mean age of 33.74 ± 0.32 Ma, with MSWD = 4.5 (Table 2; Fig. 17C). Seven εHf analyses range from −36.3 to −47.3 (average −40.1 ± 3.2).

Sample site ARRM11-12 is located on the floor of upper Lamoille Canyon 1350 m southwest of the trailhead parking area (Fig. 7A). The outcrop studied consists of a biotite monzogranite body (Snow Lake Peak type), which contains a xenolith of coarse-grained pegmatitic gneiss intruded by a leucocratic monzogranite dike (File S5). After chemical abrasion, a single zircon grain was retrieved from the pegmatitic gneiss (sample ARRM11-12c); it gave a U-Pb age of 70.4 ± 1.4 Ma. In contrast, the Snow Lake Peak–type biotite monzogranite (sample ARRM11-12a) yielded a wide range of inherited cores, including Paleoproterozoic, Mesoproterozoic, Neoproterozoic, Cambrian, and Jurassic ages (File S3). Three rim analyses yielded ages from 43.1 ± 1.0 to 37.24 ± 0.73 Ma (Table 2; File S3), εHf values from −43.6 to −23.9, Hf abundances from 7500 to 12,000 ppm, and Eu/Eu* from 0.3 to 1.6. A total of three zircon grains from the leucocratic monzogranite dike survived chemical abrasion (sample ARRM11-12b); all are CL dark (File S5) and yielded discordant U-Pb ages. As a result, an imprecise U-Pb age of 33.1 ± 2.4 Ma with MSWD of 12 was obtained (Table 2; Fig. 17D).

Sample site ARRM11-15 is located ~250 m north of eastern lake of Dollar Lakes (Fig. 7A). There, a gneissic pegmatitic dike intrudes banded granitic gneiss, and both units are intruded by a biotite monzogranite dike (File S5). Zircon grains from the biotite monzogranite (sample ARRM11-15c) consist of two types, one with fine oscillatory zoning and local patchy zoning, and the other with rounded, CL-bright inherited cores and CL-dark rims (File S5). The CL-bright inherited cores yielded discordant U-Pb Paleoproterozoic and Mesoproterozoic dates (File S3). Nine LA-MC-ICP-MS analyses resulted in a weighted-mean age of 33.90 ± 0.48 Ma, MSWD = 4.2 (Fig. 18A; Table 2), and a subset of these grains yielded an εHf value of −39.4. Hf abundances are 7500–15,500 ppm and Eu/Eu* varies from 0.2 to 0.6.

7.5 Paleogene Monzogranitic Rocks of Overlook Type

Sample WRP 08-28, a mylonitic, fine-grained, leucomonzogranitic gneiss, was collected from a road cut ~365 m south of road cuts in the granite gneiss of Thorpe Creek (Table 1). The sample yielded a myriad of SHRIMP-RG zircon ages including ca. 165, 90, 70, and 37 Ma (File S2) in addition to Archean and Proterozoic cores. Ca. 37 Ma dates were obtained from CL-brighter, magmatic, oscillatory-zoned rims (File S5) and produced a mean 206Pb/238U age of 36.93 ± 0.88 Ma (MSWD = 7.6; N = 24; Fig. 18B). Seven other analyses range between 166 and 69 Ma without forming any specific age. These older dates were obtained from mainly magmatic mantles and rims. Th/U values for spots yielding Eocene dates range from 0.03 to 0.06, with one analysis <0.01 (File S5). The combination of low Th/U and Eu/Eu* and high Hf contents in WRP 08-28 zircon suggests crystallization from anatectic melts (File S5).

A fine-grained, biotite leucomonzogranite (sample WRP 02-6) was collected along the Lamoille Canyon road near the turnoff to Camp Lamoille. This sample yielded a variety of dates (Fig. 18C) including a small cluster from zircon rims with indistinct zoning (File S5), which produced a SHRIMP-RG mean 206Pb/238U age of 32.07 ± 0.85 Ma (MSWD = 1.7; N = 6). Slightly older dates (N = 3) were obtained from other rim zones, and at least two Late Cretaceous grains were identified along with Archean and Proterozoic inherited zircons (N = 11). Spots yielding Late Cretaceous ages have variable Th/U and Eu/Eu* ratios and Hf contents, whereas spots yielding Eocene ages exhibit low Th/U, small Eu/Eu* values, and high Hf contents (File S5), indicating the likelihood that these younger zircon zones crystallized from anatectic melts. These same grains are typified by non-uniform REE patterns (File S5).

At site ARRM11-15 (Fig. 7A), the oldest rock in the outcrop is Late Cretaceous banded granitic gneiss. The banded gneiss is intruded by a coarse-grained pegmatitic gneiss, and both gneissic units are intruded by biotite monzogranite. Zircon grains from the pegmatitic gneiss (sample ARRM11-15b) typically exhibit mottled cores and oscillatory-zoned rims, but two elongate prismatic crystals lack mottled cores and are zoned from CL-darker to CL-brighter, oscillatory-zoned rims (File S5). Four concordant LA-MC-ICP-MS analyses of these latter two zircon grains yield and age of 71.01 ± 0.57 Ma, MSWD = 1.12 (Fig. 18D). A single εHf value from one of these crystals is −24.0 ± 5.4. Forty analyses of the remaining grains are concordant, with a weighted-mean age of 36.9 ± 0.24 Ma, MSWD = 7.8 (Fig. 18D). Hafnium abundances range from 15,000 to 22,000 ppm, and Eu/Eu* varies from 0.045 to 0.19. Average εHf of a subset of these grains is −27.2, the same within uncertainty as that of the Late Cretaceous grains. It is noteworthy that although zircon rims in the pegmatitic gneiss (ARRM11-15b) are of similar age to zircon in the biotite monzogranite (ARRM11-15c), their average εHf values differ by ~12 epsilon units.

Sample H5-LC-111 is a very coarse-grained, two-mica granite of the chicken-track type from a dike. It yielded a myriad of SHRIMP-RG U-Pb ages from 104 to 35 Ma (Fig. 18E) from only nine analyses. A subset of these data with ages of 57–35 Ma (Fig. 18E) from mainly very dark-CL zircon zones with pseudo-magmatic form (File S5) produced a mean 206Pb/238U age of 45.7 ± 6.4 Ma (MSWD = 236; N = 8). We think this age has no geologic meaning given that it is absent from every other sample analyzed perhaps except from data reported in Howard et al. (2011) for sample Rby-107. Because this sample has so few analyses of such disparate age results, we include it here only as another example of the challenges involved in analyzing the zircon from some Lamoille Canyon samples.

7.6 Monazite Dating

LA-MC-ICP-MS analyses of monazite from three samples of Late Cretaceous equigranular granitic gneiss typically display two age peaks (Fig. 19). Two samples (RM-26 and RD-19A) display a peak in the 88–87 Ma range and another at ca. 35 Ma, whereas sample RD-31 has two peaks of 75 and 70 Ma but no discernable Paleogene peak (Fig. 19). The 88–87 Ma monazite age of sample RM-26 is slightly younger than the U-Pb zircon age (91.8 Ma; Fig. 13A) and similar to the imprecise ca. 88 Ma zircon date determined for RD-19A (Fig. 13C). Taken together, the data indicate that monazite commonly, but not universally, retains zones whose Th-Pb dates reflect the magmatic age of the sample. Additionally, peaks at 75–70 Ma and 38–32 Ma (Fig. 19) indicate that parts of these grains were reset by later events. Analyses of monazite from gneissic biotite granite sample KW-25E display a peak at 70 Ma and at 38 Ma (Fig. 19). The 70 Ma peak is statistically indistinguishable from the U-Pb zircon age of 69.8 ± 0.6 Ma (Fig. 14A).

7.7 Zircon Inheritance in Monzogranites

Both Snow Lake Peak–and Overlook-type biotite monzogranites contain inherited zircons that range in age from Precambrian to Late Cretaceous (Fig. 20; Files S2 and S3). Snow Lake Peak–type monzogranites contain rare inherited Mesozoic zircon and an age range of Proterozoic grains (Fig. 20A), but inherited Paleozoic zircon is absent. A total of 25 Proterozoic grains were analyzed from five Snow Lake Peak–type monzogranites.

In contrast, Overlook-type monzogranites are typified by Late Cretaceous zircon inheritance as well as a sparse population of Paleozoic grains (Fig. 20B). The Late Cretaceous zircon population displays a prominent peak at ca. 70 Ma and a smaller one at ca. 80 Ma (Fig. 20C) with scant older Cretaceous inheritance. Among Overlook-type monzogranites, older inherited grains range from Archean to Neoproterozoic (Fig. 20B). A total of 48 Precambrian grains were analyzed from three Overlook-type samples.

7.8 40Ar/39Ar Data

Hornblende and biotite from metamorphic and igneous rocks were analyzed by 40Ar/39Ar incremental-release techniques summarized in File S6. Hornblende and biotite were analyzed from calc-silicate paragneiss, which constitutes much of the metamorphic wall rocks in upper Lamoille Canyon (Fig. 7A). The typical mineral assemblage of the samples is: diopside + biotite + calcic amphibole (hereafter hornblende) + scapolite + quartz + plagioclase + K-feldspar + titanite + apatite.

Hornblende from sample UL1-63 yielded an apparent age of 39.1 ± 0.2 Ma; biotite from this sample has a spectrum with a gradient in ages from ca. 18 to 0 Ma (Fig. 21A). Sample UL1-61 yielded apparent ages of 43.5 ± 0.5 Ma and 28.8 ± 0.1 Ma for hornblende and biotite, respectively (Fig. 21B). Sample RD-91 yielded apparent ages of 42.5 ± 0.2 Ma and 27.8 ± 0.2 Ma for hornblende and biotite, respectively (Fig. 21C).

The gabbro–quartz diorite suite, which is widespread in the Ruby Mountains–East Humboldt Range metamorphic core complex, has consistently yielded middle to late Eocene U-Pb (zircon) ages (ca. 40–37 Ma; e.g., Wright and Snoke, 1993; Premo et al., 2014; Zuza et al., 2021; this report). Two coarse-grained biotite-hornblende gabbros yielded complicated 40Ar/39Ar spectra. Data from hornblende from sample RL-41B are consistent with cooling through Ar closure (~500 °C) at ca. 36 Ma (in concordance with the U-Pb zircon ages), but hornblende from RL-41A, collected from the same locality, yielded a discordant spectrum, which is considered geologically meaningless (Fig. 21D). We also studied several other hornblende gabbros in other parts of the Ruby Mountains–East Humboldt Range core complex, and these samples also yielded discordant spectra. This characteristic was reported for gabbroic rocks analyzed by Dallmeyer et al. (1986). Given that the shape of the age spectra for RL-41B and RL-41A are similar, it may be that both are contaminated with excess Ar but only RL-41A has enough contamination to make the apparent Ar age greater than the corresponding U-Pb zircon age. Therefore, the interpretation that RL-41B cooled through Ar closure ca. 36 Ma must be treated with caution. The true timing of this event may have been younger than this age.

Biotite from sample UL2-24, another coarse-grained biotite-hornblende gabbro exposed in the Lamoille Canyon area, yielded a discordant age spectrum indicating excess Ar (Fig. 21E), and thus the age is geologically meaningless.

7.9 Summary and Discussion of Geochronologic Data

Initial attempts to date igneous and metamorphic events in the Ruby Mountains–East Humboldt Range core complex using K-Ar, 40Ar/39Ar, and Rb-Sr methods (e.g., Kistler et al., 1981; Dallmeyer et al., 1986) typically produced data that were confusing and, in some cases, in conflict with geological observations. Results of multi-crystal U-Pb (zircon or monazite) dating (Wright and Snoke, 1993) illustrated the complexities involved in dating these rocks in terms of inheritance and high U contents, as highlighted by later SHRIMP-RG and LA-MC-ICP-MS data (e.g., Premo et al., 2005, 2008, 2014; Howard et al., 2011; Romanoski, 2012; Moscati et al., 2023; this report). Imaging of individual zircon grains combined with isotope and trace-element analyses of distinctive crystal zones clearly indicate a complex thermal and chemical history of the core complex and permit the magmatic events to be parsed in time and integrated with core-complex development.

Spot analyses on zircon presented here and elsewhere (Howard et al., 2011) indicate the following:

  • (1)

    Cretaceous magmatism in the Ruby Mountains began with emplacement of relatively small volumes of mildly to strongly peraluminous monzogranite at ca. 91 Ma. This event is represented by the equigranular granitic gneiss of Lee et al. (2003) and the granite gneiss of Thorpe Creek. On average, these granites (ca. 91 Ma) display some of the highest εHf (zircon) values of any granites in the core complex, from −11 to −25 (Fig. 22). We now interpret the ca. 39 Ma dates on monazite in the granite gneiss of Thorpe Creek (Wright and Snoke, 1993; MacCready et al., 1997) as representing metamorphic growth or resetting during metamorphism and deformation which led to an L< S tectonite and an indication that mid-crustal, extension-related flow was occurring at ca. 39 Ma. This ductile deformation continued in the mylonitic shear zone until after the emplacement of cross-cutting dikes, at ca. 29 Ma, in the infrastructure that exhibit some of the mylonitic-zone fabric (MacCready et al., 1997).

  • (2)

    Granitic magmatism resumed from ca. 85 Ma to as young as ca. 69 Ma. U-Pb (zircon) results reported here and in Howard et al. (2011) suggest that the bulk of this magmatism spanned a narrower time range, from 75.6 to 69.1 Ma. Rocks in this group include the pegmatitic leucogranite gneiss of Lee et al. (2003) as well as banded granitic gneiss from upper Lamoille Canyon (Romanoski, 2012). If the textures and bulk compositions of dated rocks can be used as a correlation tool, this long-lived magmatism produced a batholith-scale complex of granites (e.g., Howard et al., 2011). This suite of granites is characterized by oscillatory-zoned zircon rims and individual grains with ages spanning from Late Cretaceous to as young as 35 Ma (Fig. 22). The εHf (zircon) values of this suite span a wide range, from −2 to −38 (with two more negative exceptions; Fig. 22). Moreover, the ranges of εHf values displayed by Late Cretaceous zircon in this suite are the same as εHf values of Paleogene zircon in the same samples (Fig. 22).

  • (3)

    Smaller volumes of granitic magmas were emplaced at ca. 60 Ma (Howard et al., 2011; this study). The single sample reported here is distinct from the latest Cretaceous pegmatitic gneiss and banded gneiss in having lower εHf values (most <−35) compared to the latest Cretaceous granites (most >−38; Fig. 22).

  • (4)

    A period of magmatic quiescence followed until ca. 40–39 Ma, at which time another long-lived period of magmatism began, ending at ca. 31–29 Ma. This event saw emplacement of compositionally diverse magmas, now represented by gabbro, quartz diorite, and tonalite, and more voluminous granodiorite and monzogranite. The mafic components of this Paleogene magmatism range in age from 38.8 to 36.6 Ma. These ages overlap with those of monzogranite of Snow Lake Peak type (39.9–33.1 Ma) and nearly so with those of Overlook-type granites (36.9–32.1 Ma). It is noteworthy that Cretaceous inherited zircon is rare in the mafic rocks and Snow Lake Peak–type monzogranite but abundant in Overlook-type granite (Figs. 18 and 20). The εHf values of zircon from Snow Lake Peak–type monzogranite are, with one exception, <−34. These very negative values are similar to those measured in the granitic gneiss (ca. 60 Ma) sample ARRM11-11b (Fig. 22). In contrast, εHf values of zircon from Overlook-type monzogranite are >−31 and lie within the range of εHf values displayed by the granites ranging from 75 to 69 Ma.

The thermal history of the Ruby Mountains–East Humboldt Range core complex from Late Cretaceous to middle Miocene can be summarized as follows:

  • (1)

    Emplacement of Late Cretaceous (ca. 91 Ma) granites;

  • (2)

    Emplacement of Late Cretaceous (ca. 85–69 Ma) granites and associated regional metamorphism and deformation including folding;

  • (3)

    Emplacement of Paleocene (ca. 60 Ma) granites in minor volumes, producing no metamorphism but having a different source compared to Late Cretaceous granites;

  • (4)

    Eocene to Oligocene (ca. 40–29 Ma) diverse magmatism associated with regional metamorphism, partial melting of Late Cretaceous granites, and resetting of 40Ar/39Ar ages of hornblende in the host metamorphic rocks;

  • (5)

    Cooling through 300 °C (biotite) in the early Oligocene to middle Miocene inferred to be related to the unroofing of the core complex related to crustal extension (crystal-plastic mylonitic shear zone and younger brittle detachment faulting); and

  • (6)

    Emplacement of middle Miocene (ca. 17–15.5 Ma) basalt dikes.

8.1 Data Sources

Much of the bulk-rock compositional data for Paleogene intrusive rocks is available in du Bray et al. (2007). Additional new bulk-rock major- and trace-element compositions on dated samples are presented in Table 3. File S7 (see footnote 1) presents bulk-rock Nd, Sr, and Pb isotope data and analytical methods, and File S8 presents oxygen isotope values for quartz and bulk rocks along with analytical methods.

8.2 Introduction

The analyzed Paleogene intrusive rocks display a range of SiO2 contents from ~46% to 76% and similarly wide ranges in Nd, Sr, and O isotope values (Table 3; Files S7 and S8). Among the granites, some display geochemical features similar to those of the Late Cretaceous granites whereas others are distinct in terms of trace-element and isotopic compositions. In contrast, compositions of mafic and intermediate rocks, although scattered, display narrower ranges of radioisotope compositions. These features are described below in sections 8.3 and 8.4 and their petrogenesis is evaluated in the context of the radiometric age data.

8.3 Major- and Trace-Element Abundances

The compositional range of the Paleogene monzogranites and related rocks is illustrated in Figure 23, in terms of the classification of Frost et al. (2001), and is compared to compositions of the Eocene Harrison Pass pluton (Fig. 2; Barnes et al., 2001), quartz dioritic intrusions from the East Humboldt Range, and Late Cretaceous granitic gneisses (Lee et al., 2003). Note that rocks with <55 wt% SiO2 are not plotted in this figure, but in all diagrams these more mafic rock compositions plot along extensions of the broad array of Harrison Pass pluton samples and include two gabbroic rocks dated by 40Ar/39Ar (samples RL-41A and RL-41B), which have SiO2 contents of 49.19 and 53.32 wt% (Table 3).

Distinguishing between Snow Lake Peak–type and Overlook-type granite samples (monzogranite, leucomonzogranite, pegmatite, and aplite) on the basis of major-element compositions is not possible. Therefore, general observations about these granites are made here. Granitic samples with SiO2 contents >69 wt% are nearly all peraluminous, with molar Al2O3/(CaO + Na2O + K2O) between 1.0 and 1.13 (Fig. 23A). The granites are predominantly ferroan (Fig. 23B), and they straddle the calc-alkalic to alkali-calcic boundary (Fig. 23C).

The tonalitic samples range in SiO2 content from 56% to 73%. All but one are magnesian and nearly all are calc-alkalic; they grade from metaluminous to peraluminous with increasing SiO2 and plot in the same array as mafic and main-stage rocks of the Harrison Pass pluton and quartz dioritic rocks from the southern East Humboldt Range (Figs. 23A23C). Among the Lamoille Canyon tonalites, the relatively high peraluminosity is related to the abundant biotite in these samples.

The compositions of leucocratic granitic rocks in the Harrison Pass pluton are indicated by two fields: (1) late-stage dikes and sheet-like bodies and (2) the larger, massive, two-mica monzogranite of Green Mountain Creek (Burton, 1997) (Fig. 23). The former group of granites varies widely in composition, whereas the monzogranite of Green Mountain Creek displays limited variation and is strongly peraluminous and ferroan (Figs. 23A and 23B).

Trace-element concentrations are displayed in Figure 24, in which nearly the entire range of SiO2 values is illustrated, including Paleogene gabbroic rocks, which are highlighted by a pink field. As in the major-element diagram (Fig. 23), Snow Lake Peak–type and Overlook-type granites are not distinguished from one another in Figures 24A24I because in these diagrams the two types display significant overlap. Among the granitic samples, abundances of Zr and Ba (Figs. 24A and 24C) decrease with increasing SiO2, whereas abundances of Rb, Nb, and Y and the Rb/Sr ratio increase with increasing SiO2 (Figs. 24D24G). In contrast, among the mafic and intermediate rocks, there is scant correlation between trace-element contents and SiO2 (Figs. 24A24F).

The Zr/Hf ratios of mafic and intermediate samples are uncorrelated with SiO2 and scatter broadly about the nominal mantle value of 36.3 (Fig. 24H). This ratio decreases with increasing SiO2 among the granitic rocks (Fig. 24H), which is suggestive of control by residual zircon. The Sr/Y ratios of nearly all Paleogene samples are <50 and the great majority of granitic rocks display ratios <20 (Fig. 24I). The ratios of Sr/Y (and Lu/Yb) are thought to be proxies for crustal thickness (Profeta et al., 2015; Chapman et al., 2015, 2018). As such, the low Sr/Y in our intermediate rocks may imply a moderate crustal thickness. However, it is more probable that, given an origin as crustal melts, the Sr/Y values indicate magma origins and residual mineral assemblages in the crustal source rather than the depth to the Moho.

Figure 24 allows comparison between Late Cretaceous and Paleogene granites and trace-element distinctions between Snow Lake Peak and Overlook granite types. The relationship between Sr and SiO2 contents (Fig. 24B) illustrates the following: (1) No trend exists among the mafic (gabbroic) to intermediate rocks; (2) Paleogene granites display significant overlap with Late Cretaceous (gneissic) granites, although in general the Late Cretaceous granites contain the highest Sr contents; and (3) among the Paleogene granites, there is no clear correlation between Sr and SiO2 abundance.

However, when Zr and Sr are plotted against Nb (Figs. 24J and 24K), two groups of Paleogene granites may be identified. One consists mainly of Snow Lake Peak–type granites, with relatively low Sr contents (~300 ppm) and a range of Zr contents. Within this group, Zr and Sr decrease from monzogranite to leucomonzogranite to pegmatite. This group also contains higher concentrations of Nb compared to other Paleogene granites (Figs. 24J and 24K). The other group consists mainly of Overlook-type granites, with generally lower Zr and Nb and a much larger range of Sr contents to as much as 550 ppm (Figs. 24J and 24K). The Overlook-type granites display virtually complete overlap with the Late Cretaceous granites, whereas no more than two Late Cretaceous granite compositions plot in the field of Snow Lake Peak–type granite (Figs. 24J and 24K).

8.4 Isotopic Data

Figure 25 compares Nd and Sr whole-rock isotope data and oxygen isotope data on quartz for the Lamoille Canyon Paleogene suite with Late Cretaceous granitic gneisses and with other Paleogene intrusive rocks from the core complex (Files S7 and S8, respectively; Wright and Wooden, 1991; Wright and Snoke, 1993; Barnes et al., 2001). The Paleogene rocks display an overall negative correlation of εNd and initial 87Sr/86Sr (Fig. 25A). This trend is steep for values of initial 87Sr/86Sr <0.715 and then flattens among samples with higher initial 87Sr/86Sr ratios and εNd <−20 (Fig. 25A). All gabbroic through tonalitic samples from Lamoille Canyon and all but two Harrison Pass pluton samples lie on the steep trend, as do five leucocratic monzogranite samples, whose isotopic compositions overlap those of the Late Cretaceous granitic gneisses (Figs. 25A and 25B). Monzogranite samples with εNd values <−20 are of Snow Lake Peak type, along with two samples of two-mica monzogranite of Green Mountain Creek (Harrison Pass pluton). No Late Cretaceous granitic rocks from Lamoille Canyon display such low εNd values.

Granites identified as Snow Lake Peak type on the basis of εHf, εNd, and trace-element contents also display δ18O values <+12‰, whereas granites of Overlook type and all but one Late Cretaceous granitic gneiss have δ18O values >+12‰ (Figs. 25C and 25D). Most Paleogene samples display oxygen isotope values (δ18O) of quartz <12‰, and all but one nominal Late Cretaceous granites display δ18O >+12‰ (Figs. 25C and 25D). The lower-δ18O group includes all analyzed samples from the Harrison Pass pluton, including the strongly peraluminous two-mica monzogranite of Green Mountain Creek (Fig. 25D).

8.5 Discussion of Geochemistry and Isotopic Character

Possible petrogenetic scenarios for genesis of Paleogene magmas include fractional crystallization from mafic parental magmas, assimilation combined with fractional crystallization, partial melting of crustal ± mantle rocks, and more complicated processes such as the multi-stage process that involves mantle melting, deep-crustal melting, and varying degrees of hybridization, as proposed for the Harrison Pass pluton (Barnes et al., 2001). Simple fractional crystallization cannot explain the diversity of isotopic compositions among Paleogene plutonic rocks and particularly the lack of correlation between isotope values and indices of fractionation (e.g., SiO2 content). The same argument can be made against simple assimilation plus fractional crystallization, for example, assimilation of crustal rocks in magmas of tonalitic composition. Two-end-member mixing is ruled out for the same reasons. Even multi-stage processes that involve significant lower-crustal mixing, as inferred for the Harrison Pass pluton, cannot explain the wide range of εNd and εHf displayed by Paleogene granites in Lamoille Canyon.

A successful petrogenetic model must explain the presence of the two monzogranite types, which display distinct Nd and Hf isotopic compositions, zircon inheritance, and trace-element abundances. In particular, successful models must explain the low Sr and comparatively high Nb and Zr in Snow Lake Peak–type granites as well as the correlation between Sr and Zr in this group.

In general, Overlook-type granites share trace-element and isotopic characteristics with the Late Cretaceous granites of the infrastructure: e.g., high Sr contents; low Nb contents; and nearly total overlap among Sr, Nd, Hf, and O isotope values. These data and the near-ubiquitous presence of Late Cretaceous inherited zircon in these rocks strongly suggest to us that Overlook-type magmas formed as partial melts of the Late Cretaceous granites (Howard et al., 2011). Moreover, most dated Late Cretaceous granites contain Paleogene zircon and/or monazite. These data and the presence of Overlook-type leucosomes associated with Late Cretaceous granitic gneiss (Fig. 5) indicate to us that Overlook-type granite magmas resulted from partial melting of Late Cretaceous granitic rocks of the infrastructure.

The distinctively lower δ18O, εNd, and εHf of Snow Lake Peak–type granites compared to the Late Cretaceous granites means that Snow Lake Peak–type granitic magmas cannot be partial melts of the Late Cretaceous granites. Instead, we suggest that Snow Lake Peak–type monzogranite is best explained by partial melting of deep, old, crustal rocks. The decrease in Sr, Ba, and Zr contents with increasing differentiation (e.g., SiO2 content; Fig. 24) and low Sr/Y indicate the presence of residual feldspar and zircon in the source (see also Lee and Barnes, 1997). These features are distinct from those of the source of Late Cretaceous granitic magmas, which were interpreted to contain residual garnet (Lee et al., 2003). The Snow Lake Peak–type granites share trace-element and isotopic features with the peraluminous monzogranite of Green Mountain Creek in the Harrison Pass pluton (Figs. 24 and 25), which was interpreted to be a partial melt of old, lower-crustal rocks (Barnes et al., 2001). We suggest that the same explanation applies to Snow Lake Peak–type granites. These low εNd and εHf values are consistent with Neoarchean to Paleoproterozoic sources of the Snow Lake Peak–type granites.

Petrogenesis of the gabbroic, dioritic, and tonalitic rocks is complicated because when taken as a group, these rocks lack geochemical signatures of a petrologic process (Figs. 23 and 24) and none display isotopic signatures of a normal or depleted mantle source (Fig. 25). It is possible to interpret the Nd and Sr isotopic data as arising from partial melting of enriched mantle. However, this model does not explain the high δ18O values (seen in main-stage rocks of the Harrison Pass pluton; Barnes et al., 2001). Instead, we suggest that the isotopic features of this group of rocks reflect crustal assimilation of partial melts of crustal rocks in mafic, mantle-derived magmas. The origin of the crustal melts potentially includes melts of Late Cretaceous granites in levels below the middle crust.

Thus, we find it likely that trace-element and isotopic variation among the Paleogene plutonic granitic rocks reflects: (1) assimilation of deep-crustal rocks by mantle-derived mafic magmas to form the wide range of mafic to intermediate magmas now seen as gabbroic through tonalitic rocks, (2) deep-crustal melting of old crustal rocks to form the Snow Lake Peak–type magmas, and (3) local partial melting of Late Cretaceous granites to form the Overlook-type magmas. This complex interplay of crustal melting and crustal assimilation powered by deep-crustal emplacement of mafic magmas is similar to that proposed for the Harrison Pass pluton (Barnes et al., 2001) and intrusions in the Snake Range and Egan Range complexes of Nevada (Grunder, 1992).

Estimates of temperature (T) and pressure (P) of emplacement of plutonic rocks are problematic because a range of conditions may be recorded during cooling. The presence of coexisting plagioclase and calcic amphibole in some samples of tonalite and biotite monzogranite permits estimates of P and T. The Al-in-hornblende barometer may be used when the buffer assemblage quartz + alkali feldspar + plagioclase + hornblende + biotite + iron-titanium oxide + titanite is present (Hammarstrom and Zen, 1986; Hollister et al., 1987). Among our samples, three tonalitic dikes, a single biotite monzogranite, and three leucocratic granites with amphibole clusters contain some or all of the required equilibrium assemblage, although samples commonly lack titanite and some tonalites lack alkali feldspar. However, Anderson and Smith (1995) indicated that the presence or absence of these minerals does not affect P estimates, and our results show no distinction between samples with and without these minerals. Al-in-hornblende pressure estimates were obtained employing the calibration of Anderson and Smith (1995). This calibration incorporates the effect of T as determined by the hornblende-plagioclase geothermometers of Holland and Blundy (1994). These two geothermometers use edenite-tremolite and edenite-richterite exchanges; Anderson and Smith (1995) preferred results of the latter. In addition, Anderson and Smith (1995) showed that Fe-rich calcic amphibole compositions are outside the limits of experimental studies and tend to yield excessively high P estimates. Finally, because the Anderson and Smith (1995) calibration uses amphibole and plagioclase rim compositions, it yields near-solidus to solidus P and T estimates.

Three tonalite samples contain amphibole with Fe/(Fe + Mg) values within the range Anderson and Smith (1995) viewed as appropriate for Al-in-hornblende barometry. Amphibole from monzogranite samples is too Fe rich. P estimates from the tonalite samples range from 520 to 590 MPa with individual uncertainties of ±60 MPa, and T estimates range from 764 to 661 °C using the edenite-tremolite thermometer.

Putirka (2016) proposed an amphibole geothermometer that is not dependent on the equilibrium assemblage or pressure. Using this algorithm yielded T estimates for amphibole in the tonalitic rocks of 815–753 °C and for amphibole in monzogranites of 778–724 °C. Published calibration uncertainties are ±30 °C and the standard deviation about mean T values ranges 5–31 °C. Higgins et al. (2022) presented an amphibole thermobarometer-chemometer calibrated using random forest machine learning. P estimates for tonalitic and monzogranitic samples ranged widely (400 ± 10 to 930 ± 100 MPa, n = 7), with four samples yielding estimates from 400 ± 10 to 440 ± 17 MPa. Temperature estimates varied from 822 to 849 °C with typical individual uncertainties of ~15 °C.

Because the majority of monzogranite samples are medium to fine grained and equigranular, use of the zircon saturation thermometer may be appropriate. Use of the Watson and Harrison (1983) algorithm indicates zircon saturation among monzogranite samples was mainly from ~830 °C to 700 °C, with estimated T decreasing with increasing bulk-rock SiO2 contents. Although these T estimates may clearly be affected by presence of inherited zircon, they do encompass the range of amphibole T estimates using Putirka’s (2016) geothermometer and overlap with results of the Higgins et al. (2022) geothermometer.

Finally, the liquidus temperature of a typical biotite monzogranite composition was calculated using rhyolite–MELTS (Gualda et al., 2012; Gualda and Ghiorso, 2015). Calculations were run at 550 MPa using the quartz-magnetite-fayalite oxygen buffer and H2O contents from 3 to 5 wt%. With increasing H2O content, liquidus T decreased from 895 to 860 °C.

Thus, we interpret the P of emplacement to be ~400–550 MPa. The combination of rhyolite–MELTS calculations with hornblende and zircon saturation thermometry indicates that the T of monzogranitic magmas could have been as high as 890 °C and clearly reached or exceeded T of ~830 °C. The estimated emplacement P of the biotite monzogranite of 400–550 MPa would correspond to emplacement depths of ~16–22 km if lithostatic.

Intrusive dikes offer opportunities for analysis of paleostresses (Nakamura, 1977; Zoback et al., 1981; Zoback, 1989; Ildefonse et al., 1993). Dikes generally intrude by processes akin to hydrofracturing, in which fluid overpressure leads to tensile failure perpendicular to the least principal compressive stress. Unless they show a clear influence by pre-existing structures, dikes tend to map the orientation of the plane formed by the two greater principal stresses (Anderson, 1951; Pollard, 1987). Spatial or sequential changes in dike orientation thus can be used as a guide to track paleostresses.

The tabular nature of many of the dikes and sheets measured for this study, excepting the irregularly shaped tonalitic dikes and leucosome-like Overlook-type granite bodies, therefore, are interpreted to track the local stress. We envisage a host-rock environment of slow plastic yielding, with rheology that varies with rock type.

Temperature and pore pressure at the level of the crust intruded by the Snow Lake Peak intrusion likely combined to lower strength and inhibit the maintenance of much differential stress. The tendency for the dikes in upper Lamoille Canyon to form crudely orthogonal orientation arrays (Fig. 6) suggests to us that the least principal stress may have flipped back and forth in orientation direction as fluid or dike pressure enhanced or relieved the relative magnitudes of the principal stresses (cf. McCarthy and Thompson, 1988; Goodwin et al., 1989). Dikes may record relatively constant orientation for the three principal stress directions during the period of intrusion. However, fluctuations in their relative magnitudes could allow the direction of least principal stress to switch back and forth among the three. Assuming small differential stress, small fluctuations in the magnitudes of principal stresses could lead to interchange of their relative magnitudes and to successive dike emplacement in each of the principal directions. Such fluctuations potentially could be driven with each new intrusion as the addition of rock bulk against the local least principal stress direction increased confinement of that direction (McCarthy and Thompson, 1988; Goodwin et al., 1989). The orientation data in upper Lamoille Canyon are thus consistent with crudely constant orientation for the principal stress directions during dike emplacement, low deviatoric stress, and fluctuation of the relative stress magnitudes. The state of stress probably fluctuated along with fluid-pressure changes.

The half-rainbow shapes of three biotite monzogranite dikes (Fig. 26A) hold further clues to stress conditions during the time of intrusion. These parallel dikes dip 45°–50° east (toward 070° azimuth) at a gently west-dipping stratigraphically inverted contact of metaquartzite over the subjacent marble unit. In the metaquartzite above that contact, the dikes curve to horizontal (Fig. 26A). If considered as stress trajectories, the rainbow dikes indicate that stress orientations in the quartzite depended on the proximity to the metaquartzite-marble contact. High shear stress is indicated for the metaquartzite-marble contact, inasmuch as directions of maximum shear stress are 45° from principal stress directions. We infer that the stresses in the metaquartzite relate to slow plastic yielding and that the basal contact served as a rough frictional boundary (cf. Hubbert and Rubey, 1959). Pre-existing intricate folding and a concentration of irregular pre-Tertiary intrusions of pegmatitic leucogranitic gneiss, observed at this boundary, may have served to inhibit shear strain and to build up shear stress at the boundary. Shear stress at this boundary may also have built up due to differential plastic yielding of the quartzite versus marble-plus-gneiss host rocks. The curving stress trajectories in the quartzite resemble those calculated for an eastward-sliding thrust block or for compressive eastward flow in a glacier (Hafner, 1951; Hubbert, 1951; Nye, 1952, 1967). The direction of curvature of the dikes implies compressive stresses at the time of intrusion that allowed upward and eastward expansion within the quartzite.

Paradoxically, this horizontal compression and implied east-over-west yielding exactly opposes the strain deduced for the Oligocene extensional mylonitic shear zone carapace, which may be similar in age and projects <1 km above the rainbow dikes (Howard, 2000). That mylonitic zone instead records crustal attenuation and top-to-the-west extensional shear. If the mylonitic extensional shearing operated at ca. 29 Ma when the rainbow dikes were emplaced, the stress conditions in their host rocks possibly relate to a tendency for isostatic return flow toward the area being unroofed (cf. Block and Royden, 1990; MacCready et al., 1997).

11.1 Relationship of the Upper Tier of the Core Complex to Extension at Higher Crustal Levels

The upper structural tier of the Ruby Mountains–East Humboldt Range metamorphic core complex is a cover terrain consisting of unmetamorphosed middle Paleozoic to Triassic sedimentary rocks unconformably overlain by middle Eocene volcanic and volcanogenic strata. These rocks form the hanging wall of the Ruby Mountains–East Humboldt Range low-angle, normal-sense detachment fault (Fig. 2). This low-angle fault, therefore, must be younger than 40Ar/39Ar-dated middle Eocene volcanic rocks in the southeastern East Humboldt Range (Brooks et al., 1995a, 1995b). In the Secret Creek gorge area ~32 km NNE of the upper Lamoille Canyon area, the detachment-fault system is overlain unconformably by middle Miocene rocks of the Humboldt Formation and rhyolite flows (ca. 15 Ma) (Dee et al., 2015; Snoke et al., 2021). However, a low-angle normal fault truncates a basalt dike (ca. 17 Ma) in the nearby Tent Mountain quadrangle (Snoke, 1980; Snoke et al., 2021; Zuza et al., 2021). Therefore, in that area, final slip along the detachment-fault system is bracketed between 17 and 15 Ma. In the southern Ruby Mountains, Colgan et al. (2010) determined slip along the Ruby Mountains detachment fault as 17–15 and 12–10 Ma, which led to rapid unroofing of the southern Ruby Mountains.

Newer 40Ar/39Ar and (U-Th)/He studies indicate an episodic cooling history in the Ruby Mountains–East Humboldt Range core complex (McGrew et al., 2019; McGrew and Metcalf, 2021). These studies record a late Eocene to Oligocene cooling event in the East Humboldt Range inferred to be related to extensional exhumation. Extensional cooling apparently slowed in the middle Oligocene, restarted in the latest Oligocene to early Miocene, again slowed between ca. 21 Ma and ca. 17.5 Ma, and subsequently restarted in the middle Miocene and lasted until ca. 10 Ma (McGrew and Metcalf, 2021). Basin and Range normal faulting began after ca. 10 Ma and continues to the present (Colgan and Henry, 2009). Rapid denudation and uplift are recorded by quenching through 300 °C to ~100 °C at ca. 23 Ma (Dokka et al., 1986; Dallmeyer et al., 1986; McGrew and Snee, 1994). By 17–15.5 Ma, the core complex was denuded to depths shallow enough that gas could exsolve from basalt magma as indicated by vesicles in dikes.

West of the Ruby Mountains, a lacustrine system existed during the middle Eocene, manifested by the Elko Formation (Solomon et al., 1979; Canada et al., 2020). As indicated by paleoflora (Wolfe et al., 1998) and stable isotopes (Sewall and Fricke, 2013; Snell et al., 2014; Cassel et al., 2014), this formation was deposited at high elevations upon the erosional highland of the Nevadaplano (DeCelles, 2004), although compare to Lund-Snee et al. (2016). Deposition in the Elko Formation abruptly ended with the eruption of voluminous ash-flow tuffs at 41.5 and 40.1 Ma (Canada et al., 2020). Use of (U-Th)/(He-Pb) double dating of short-lag-time detrital zircons from the Elko Formation indicates rapid unroofing of the adjacent Ruby Mountains–East Humboldt Range core complex in the middle Eocene (42–40 Ma) (Canada et al., 2020). The Oligocene (ca. 29.6 Ma) tuff of Campbell Creek, erupted from a caldera in north-central Nevada, spread eastward in paleovalleys cut into the Nevadaplano as far as the East Humboldt Range (Henry et al., 2012; McGrew and Snoke, 2015). The interpreted thickness of the crust during the formation of the Nevadaplano reached ≥55 km (Chapman et al., 2015), with the lower crust consisting of igneous and metamorphic rocks that developed during Late Cretaceous tectonic thickening (Coney and Harms, 1984; McGrew et al., 2000; Hallett and Spear, 2014, 2015). Destabilization of the thickened crust of the Nevadaplano, leading to its collapse and thinning, has been interpreted as related to the progressive delamination of the Farallon slab and upwelling of the underlying asthenosphere (Humphreys, 1995; Canada et al., 2020). We relate late Eocene to Oligocene denudational cooling of the Ruby Mountains–East Humboldt Range core complex to this crustal thinning and emphasize its coincidence with the plutonism and regional northeastern Nevada volcanism that occurred ca. 43–29 Ma (Figs. 7 and 8). Upwelling asthenosphere likely drove mantle-derived melts to pond in the lower crust where assimilation produced the gabbroic magmas and transferred heat to the crust, resulting in partial melting of lower-crustal rocks to yield Snow Lake Peak–type granitic magmas and of middle-crustal granites to yield Overlook-type granitic magmas. The emplacement, ponding, and differentiation of granitic and mafic magmas in the lower crust as well as assimilation and partial melting may have been related to this tectonic event. In northeastern Nevada, this magmatism began during the middle Eocene (ca. 43–40 Ma) and lasted to the early Oligocene, producing volcanic rocks and plutonic rocks such as the Snow Lake Peak and Castle Lake intrusions (Figs. 7 and 8).

11.2 Relationship of the Lower Tier of the Core Complex to Ductile Deformation

The nearly orthogonal inferred north-south and east-west brittle extension recorded by the pegmatite and other dikes in upper Lamoille Canyon may relate to the nearly orthogonal elongation and mineral lineations found in the country rocks and Tertiary intrusions north of Lamoille Canyon discussed by MacCready et al. (1997). These authors proposed that north-south-trending lineation in rocks in the infrastructure record plastic north-south shear flow in the middle crust. Brittle north-south extension is expressed by steep, east-west-striking dikes. North-south-striking dikes in upper Lamoille Canyon record brittle extension that possibly relates to east-west-trending stretching lineations in the mylonitic carapace exposed ~8 km to the west. Those lineations also record high-temperature (amphibolite-facies) plastic shear in the azimuth of extension. Deformation in this shear zone occurred between ca. 29 and 23 Ma (Dallmeyer et al., 1986; Dokka et al., 1986; Wright and Snoke, 1993; MacCready et al., 1997). However, magmatism and plastic flow in the infrastructure ended in the early Oligocene (ca. 29 Ma) as indicated by cross-cutting biotite monzogranite dikes, which truncate foliation in the wall rocks (Fig. 26B). Surprisingly, in light of the east-west extensional shearing, east-west-striking dikes dominate over the north-south-striking dikes. Furthermore, the presence of subhorizontal dikes indicates some degree of vertical inflation of the section during intrusion, not a result expected in an environment of east-west extension but perhaps expected in an environment of magma inflation during extension.

A buoyant-gneiss-dome model for Cordilleran core complexes has been applied to the Ruby Mountains–East Humboldt Range (Howard, 1980; Zuza et al., 2022; Levy et al., 2023; Zuza and Dee, 2023) and Albion–Raft River–Grouse Creek core complexes (Konstantinou et al., 2012, 2013). Other studies of Cordilleran core complexes such as the Anaconda complex in western Montana (Howlett et al., 2021) and Catalina complex in southeastern Arizona (Ducea et al., 2020) also indicate that magmatism was essential in core-complex development. Magmatic intrusion has been suggested as the driving force for active metamorphic core complexes in eastern Papua New Guinea (Hill et al., 1995). Therefore, the upward vertical flow of magmatic and/or migmatitic rocks and the synchronous development of marginal shear zones is a viable model to explain many aspects of a metamorphic core complex (Whitney et al., 2004b). Levy et al. (2023), Zuza and Cao (2023), and Zuza and Dee (2023) contrasted buoyant doming with the rolling-hinge model. They highlighted general shear in the mylonitic shear zone and a significant gap in the time of the formation of the shear zone and detachment faulting. Although they favored the buoyant-doming model, they noted an alternative model is a hybrid extensional gneiss dome with a marginal rooted mylonitic shear zone (Levy et al., 2023, their figure 1C). That is the model that we favor for the origin of the Ruby Mountains–East Humboldt Range core complex (Fig. 27). The upward flow of the core complex was facilitated by the widespread emplacement of the Paleogene granitic and gabbroic to quartz dioritic magmas as small intrusions, sheets, and dikes as well as partial melting of Late Cretaceous granites presently exposed and deeper in the middle crust. Coupled with the magmatism was the role of convergent crustal flow in the infrastructure as hypothesized by MacCready et al. (1997) and Litherland and Klemperer (2017). Structurally above the infrastructure, shear strain and stretching in the 200–300-m-thick mylonitic shear zone was ongoing from ca. 29 to 23 Ma following the magmatism. Numerous sense-of-shear microstructures indicate a west-rooted zone of progressive simple shear (Lister and Snoke, 1984). There is no evidence of an antithetic shear zone in the core complex as predicted in the buoyant-gneiss-dome model. Middle Eocene to early Oligocene intrusive rocks were overprinted by the mylonitic shear zone during exhumation of the core complex. An upper-crustal, brittle middle Miocene detachment fault captured the mylonitic shear zone, and Miocene unroofing produced thick sedimentary accumulations on the flanks (e.g., Humboldt Formation). Late Miocene to Holocene Basin and Range normal faults cut the detachment fault (Snoke et al., 2021; Zuza and Dee, 2023).

  • (1)

    Middle to late Eocene to early Oligocene magmatism in the upper Lamoille Canyon region of the Ruby Mountains encompassed a range of compositions from gabbro through leucomonzogranite.

  • (2)

    Aside from the late Eocene Harrison Pass pluton ~34 km SSW, individual middle Paleogene intrusions in the Ruby Mountains are small and many are sheet like. Emplacement was in the range of 400–550 MPa, which if lithostatic would correspond to depths of 16–22 km.

  • (3)

    Small-volume Eocene gabbros signal advection of mantle heat as an important driver of the magmatism. The εNd of mafic rocks may indicate derivation from old, enriched, lithospheric mantle, but high δ18O values of similar rocks in the Harrison Pass pluton argue for significant contamination by crustal rocks or melts.

  • (4)

    Hf isotope values in zircon indicate that the sources of biotite monzogranites (Snow Lake Peak type) were old (Neoarchean to Paleoproterozoic) crustal rocks.

  • (5)

    Small leucogranite bodies ca. 39–29 Ma have εHf (zircon) identical to the εHf of country-rock Late Cretaceous pegmatitic leucogranite orthogneiss as well as of leucosomes associated with Late Cretaceous pegmatitic leucogranite (Overlook type). We interpret these features to indicate that the voluminous Late Cretaceous pegmatitic leucogranite, which constitutes the greatest bulk of the infrastructure, underwent partial melting ca. 39–29 Ma.

  • (6)

    The small Paleogene intrusions clearly advected heat to the middle crust, but the thermal event was even more pervasive, given that it produced persistent Eocene to Oligocene (ca. 39–32 Ma) zircon in Late Cretaceous host granites, and it partially melted Late Cretaceous granites to produce Overlook-type Paleogene leucogranite.

  • (7)

    Widespread middle-crustal Eocene anatexis left the infrastructure weakened, locally as a mobile crystal mush. By the early Oligocene (ca. 29 Ma), the solidifying infrastructure was intruded by cross-cutting biotite monzogranite dikes. Some of these dikes were mylonitized in the crystal-plastic shear zone that formed above the infrastructure.

  • (8)

    Addition of mafic magma in the lower crust, transfer of granitic magmas into the middle crust, and consequent partial melting in the middle crust predisposed the crust for plastic extensional flow. Crustal softening was potentially a cause of extensional collapse of the Nevadaplano.

  • (9)

    We favor a hybrid model of formation of the core complex that combines diapiric rise of mobile intrusion-softened crust and Paleogene unidirectional extensional shearing in the course of collapse of the softened crust.

1Supplemental Material. File S1: Mineral analyses (amphibole, biotite, and plagioclase). File S2: SHRIMP-RG U-Pb data. File S3: LA-MC-ICP-MS age data and LA-Q-ICP-MS trace-element data with methods. File S4: LA-MC-ICP-MS Hf isotope (zircon) data. File S5: Supporting data and images for zircon geochronology. File S6: 40Ar/39Ar data on hornblende and biotite. File S7: Bulk-rock Sr, Nd, and Pb isotope data. File S8: Oxygen isotope values of quartz and bulk rocks with methods. Whole-rock chemical analyses of rocks of the middle to late Eocene and early Oligocene monzogranitic suite. Refer to U.S. Geological Survey Data Series 244 (du Bray et al., 2007). Please visit https://doi.org/10.1130/GEOS.S.24915717 to access the supplemental material, and contact editing@geosociety.org with any questions.
Science Editor: Andrea Hampel
Associate Editor: Michael L. Williams

Many geologists have visited Lamoille Canyon with us on field trips and offered valuable comments on the field and geochronological, geochemical, and isotopic data. We thank Phyllis Ranz for the preparation of many figures in this paper. We thank Richard Allmendinger for making the Stereonet 11.0.9 software package freely available to us and the entire geological community (Allmendinger et al., 2013). We thank Joe Wooden for providing SHRIMP data for a gabbro sample. Joe Colgan helped in the collection of several samples for SHRIMP analysis. Jim Wright first recognized an Eocene monazite age for a sample of leucomonzogranite in upper Lamoille Canyon. We thank Jamey Jones, Michael Wells, two anonymous Geosphere reviewers, and Associate Editor Michael Williams for helpful and constructive suggestions improving the manuscript. We also thank Nancy Stamm and Geneva Chong, U.S. Geological Survey editors, for their edits and comments, which improved our paper. This research was partially funded by U.S. National Science Foundation grants EAR-9627958 to Snoke and EAR-9627814 to Barnes. Any use of trade, firm, or product names is for descriptive purposes only and does not imply endorsement by the U.S. Government.

Gold Open Access: This paper is published under the terms of the CC-BY-NC license.